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Thesis
Geochronology, geochemistry, and isotopic composition (Sr, Nd, Pb)
of Tertiary porphyry systems in Ecuador
SCHUTTE, Philip
Abstract
This thesis presents geochronologic, geochemical, and isotopic data on Tertiary arc magmas
associated with porphyry-related ore deposits in the northern Andes of Ecuador. In detail,
ages on host and porphyry intrusion emplacement (U-Pb zircon) and hydrothermal
alteration/mineralization (U-Pb titanite, Re-Os molybdenite) were obtained, combined with
whole-rock geochemical analysis of least altered igneous rock samples. These data are
discussed in a regional frame with respect to the geodynamic context. The following main
conclusions apply: (1) porphyry-related ore deposits in Ecuador are mainly Miocene in age,
and represent the northern extension of the Miocene metallogenic belt of northern-central
Peru; (2) intrusions associated with mineralization are compositionally highly variable; a
systematic association of mineralization with a specific trace element signature cannot be
observed; (3) flat slab settings produce favorable exposure conditions for porphyry-related ore
deposits; a systematic association of ore deposit formation and seamount chain (Carnegie
Ridge) subduction is not observed.
Reference
SCHUTTE, Philip. Geochronology, geochemistry, and isotopic composition (Sr, Nd, Pb)
of Tertiary porphyry systems in Ecuador. Thèse de doctorat : Univ. Genève, 2009, no. Sc.
4166
URN : urn:nbn:ch:unige-63675
Available at:
http://archive-ouverte.unige.ch/unige:6367
Disclaimer: layout of this document may differ from the published version.
[ Downloaded 19/10/2016 at 16:25:00 ]
UNIVERSITÉ DE GENÈVE FACULTÉ DES SCIENCES Département de Minéralogie Prof. Urs Schaltegger Dr. Massimo Chiaradia Geochronology, Geochemistry, and Isotopic Composition (Sr, Nd, Pb) of Tertiary Porphyry Systems in Ecuador THÈSE présentée à la faculté des sciences de l'Université de Genève pour obtenir le grade de Docteur ès sciences, mention Sciences de la Terre par Philip SCHÜTTE de Allemagne GENÈVE 2010
TABLE OF CONTENTS Acknowledgements ...........................................................................................................................................iii
Abstract ..............................................................................................................................................................v
Resumen ...........................................................................................................................................................vii
Résumé étendu..................................................................................................................................................xi
CHAPTER I – THEORETICAL BACKGROUND AND AIMS OF THESIS Introduction....................................................................................................................................................... 1
References ......................................................................................................................................................... 3
CHAPTER II – GEODYNAMIC CONTROLS ON TERTIARY ARC MAGMATISM IN ECUADOR: CONSTRAINTS FROM U‐PB ZIRCON GEOCHRONOLOGY OF OLIGOCENE‐MIOCENE INTRUSIONS AND REGIONAL AGE DISTRIBUTION TRENDS Abstract ............................................................................................................................................................. 5
Introduction....................................................................................................................................................... 6
Regional Geology of Ecuador ............................................................................................................................ 6
Tertiary arc segmentation and geology of investigated magmatic centers...................................................... 8
Analytical techniques and sample material .................................................................................................... 14
Results ............................................................................................................................................................. 18
Discussion ........................................................................................................................................................ 25
Conclusions...................................................................................................................................................... 34
References ....................................................................................................................................................... 34
Appendix I – Accommodation of convergence obliquity at the Ecuadorian margin throughout the Tertiary ...................................................................................................................................................... 39
Appendix II – Overriding plate structural controls on the spatio‐temporal distribution of Tertiary plutons in Ecuador........................................................................................................................................... 42
Appendix III – Accuracy of published K‐Ar (and ZFT) ages of Tertiary intrusions in Ecuador.......................... 48
Appendix IV – Data tables ............................................................................................................................... 54
CHAPTER III ‐ THE MIOCENE METALLOGENIC BELT OF ECUADOR: CONSTRAINTS FROM NEW RE‐OS MOLYBDENITE AND U‐PB TITANITE AGES OF PORPHYRY‐RELATED ORE DEPOSITS Abstract ........................................................................................................................................................... 65
Introduction..................................................................................................................................................... 66
Regional geology and geodynamic setting...................................................................................................... 68
Local geology of Miocene Ecuadorian ore deposits investigated in this study............................................... 69
Sampling and analytical techniques ................................................................................................................ 73
Results ............................................................................................................................................................. 74
Discussion ........................................................................................................................................................ 76
Conclusions...................................................................................................................................................... 92
References ....................................................................................................................................................... 92
i
CHAPTER IV ‐ CRUSTAL BASEMENT ARCHITECTURE IN ECUADOR EXPLORED BY SR, ND, AND PB ISOTOPIC COMPOSITIONS OF TERTIARY‐QUATERNARY ARC MAGMAS Abstract ........................................................................................................................................................... 97
Introduction..................................................................................................................................................... 98
Geological framework ................................................................................................................................... 100
Methodology ................................................................................................................................................. 100
Results ........................................................................................................................................................... 103
Discussion ...................................................................................................................................................... 106
Conclusions.................................................................................................................................................... 112
References ..................................................................................................................................................... 112
Appendix: Data tables ................................................................................................................................... 115
CHAPTER V ‐ ADAKITE‐LIKE FEATURES IN LATE OLIGOCENE TO LATE MIOCENE ECUADORIAN ARC MAGMAS AND THEIR SIGNIFICANCE FOR PORPHYRY‐RELATED ORE DEPOSITS Abstract ......................................................................................................................................................... 119
Introduction................................................................................................................................................... 120
Tertiary‐Quaternary adakite‐like magmatism in Ecuador............................................................................. 120
Regional geology and geodynamic setting.................................................................................................... 122
Sampling and analytical techniques .............................................................................................................. 126
Results ........................................................................................................................................................... 128
Rare earth element distribution patterns ..................................................................................................... 134
Adakite‐like features of Late Tertiary Ecuadorian arc magmas .................................................................... 139
Isotopic constraints on shallow vs. deep crustal magma evolution.............................................................. 144
Significance of adakite‐like features for Late Oligocene to Late Miocene porphyry‐related mineralization in Ecuador.............................................................................................................................. 151
Conclusions.................................................................................................................................................... 152
References ..................................................................................................................................................... 153
Appendix I – Rock alteration and element mobility in porphyry‐related hydrothermal systems ................ 157
Appendix II – Data tables............................................................................................................................... 159
CHAPTER VI ‐ GENERAL CONCLUSIONS AND OUTLOOK Conclusions.................................................................................................................................................... 179
Outlook .......................................................................................................................................................... 180
ii Acknowledgements Big thanks to Massimo Chiaradia for initiating this project, for his continuous and patient support through‐
out its lifetime (both in the lab and in the office), and for doing a superb job with constructive and critical thesis reviewing which was highly appreciated. Successful field work in Ecuador would have been impossi‐
ble without the organizational support of Bernardo Beate whose personal commitment to this project is gratefully acknowledged. Urs Schaltegger, Kalin Kouzmanov, and Othmar Müntener are thanked for accept‐
ing to be part of the committee of this PhD project, and for valuable discussions and critical comments re‐
garding various aspects of the material presented in this thesis during all stages of its compilation over the last couple of years. Urs' administrative support is also gratefully acknowledged. Many individuals at the Universities of Geneva, Lausanne, and Arizona, contributed to the success of this project. As far as Geneva is concerned, my knowledge of the geology of Ecuador improved from valuable discussions with Richard Spikings, Miguel "No More Mr. Nice Guy" Ponce, and, in particular, Diego Vil‐
lagómez, who are all gratefully acknowledged. I also thank Aldo Bendezú for good discussions and for pos‐
ing critical questions which stimulated me to look at geological problems of all sorts from a different per‐
spective, and to think again on many aspects of metallogenesis, magmatism, and tectonics. Big thanks to Maria Ovtcharova, Blair Schoene, and Urs for incredible support in the Geneva zircon lab, with mass spectrometric analysis, and with U‐Pb data reduction. I further thank Fabio Capponi for XRF analysis and Jean‐Marie Boccard for thin section preparation; I had quite a lot of samples for both of them, and they always took it on good‐humored and provided excellent results which were highly appreciated. Thanks for analytical support are further due to Alex Ulianov (MC‐ICP‐MS, Lausanne) and Rosanna Martini (SEM, Ge‐
neva). At the University of Arizona, Fernando Barra and Victor Valencia did a great job in providing Re‐Os molybdenite and U‐Pb zircon analyses for this study, and both are gratefully acknowledged for their valu‐
able contributions. Many administrative aspects of this project were taken care of by Sofia Saldana, Jacque‐
line Berthoud, and Ursula Eigenmann, and I warmly thank each of them. I gratefully acknowledge SEG (Society of Economic Geologists), EDSM (Ecole Doctorale en Sciences des Mi‐
néraux), and the Bourse Lombard for providing travel grants to support congress participation and field‐
work of this project. Thanks in particular to Mike Dungan for investing a lot of time and personal effort into EDSM and for organizing a number of top‐level short courses, as well as for repeated tech support bribery. Furthermore, I am grateful to Lluís Fontboté for encouraging me to participate in several excellent courses and excursions on ore deposits whose informative value was priceless. Many students at Geneva University provided help and advice throughout all stages of this project. In par‐
ticular, I thank Diego Villagómez, Jenny Skoog, Thierry Bineli Betsi, and Régine Baumgartner for lab support in the beginning, Léo Mastrodicasa for advice during the initial stages of computer administration, and Toufik Bekaddour and Diego Villagómez for help with editing the French and Spanish abstracts which was highly appreciated. This project benefited substantially from countless helpful people and exploration companies in Ecuador whose enthusiastic support with logistics and knowledge was key to its success, particularly in a politically sometimes challenging climate. These include (with their affiliations back then): Patricio Salazar (Escuela Politécnica Nacional & Ascendant Copper), Carlos Moncayo and James Stonehouse (Ascendant Copper), Luis Bravo and Edgar Almeida (Dynasty Metals & Mining), Francisco Soria and Miguel Ponce (IMC), Luis Lucero and Patricio Perez Salazar (Iamgold), John Bolaños and Danilo Ortega (Ecuagold), Graeme Smith, Don Allen, and Eduardo Vaca (Atlas Moly), Osman Poma (Channel Resources), and Christian Vallejo (Curimining); apologies for any missing names here. Thanks are also due to the Swiss Embassy in Quito for providing quick and efficient help with sample export, and the Ministerio de Energía y Minas del Ecuador (especially Luis Pilatasig and Fabiola Alcocer) for providing otherwise inaccessible maps and literature. Last but not least, I thank my family for their continuous support in all these years. iii
iv Abstract This Ph.D. thesis presents geochronologic data on Late Tertiary arc magmatism associated with porphyry‐
related ore deposits (Chapter 2), and the timing of intrusion‐related mineralization and alteration in Ecua‐
dor (Chapter 3), as well as on the isotopic (Chapter 4) and geochemical (Chapter 5) composition of Oligo‐
cene‐Miocene arc volcanic and plutonic rocks spatially associated with porphyry systems. Chapter‐specific abstracts precede each part of the thesis and should be consulted for more detailed information on each topic. A general introduction (Chapter 1) explains some of the ideas this project was originally based on, and general conclusions joining its individual parts can be found in Chapter 6. Ecuador hosts a large number of Tertiary arc‐related granitoids and volcanics whose geochronologic char‐
acterization is mostly based on K‐Ar (and zircon fission track) data. Chapter 2 presents the first regional‐
scale dataset of U‐Pb zircon ages for multiple plutons and porphyry intrusions of the Ecuadorian Western Cordillera and southern Sierra region, allowing a robust time calibration of the Tertiary intrusive history of Ecuador. Zircons were dated either by means of isotope dilution thermal ionization mass spectrometry (TIMS) or by laser ablation multi‐collector inductively coupled plasma mass spectrometry (LA‐MC‐ICP‐MS). Except for a single sample, all investigated intrusions completely lack externally inherited zircons suggesting a dominantly zircon‐poor oceanic basement; antecrystic zircon components are relatively abundant in sev‐
eral plutons indicating that recycling of older arc granitoids took place. Where both K‐Ar and U‐Pb data ex‐
ist for a given intrusive system, ages obtained by the different methods are usually concordant within 1‐4 m.y. implying that K‐Ar data may be used as a semi‐accurate proxy for Tertiary arc magmatism in Ecuador on a regional scale. Spatio‐temporal distribution trends of Tertiary intrusions and arc volcanics indicate a Late Oligocene to Early/Mid‐Miocene arc magmatic flare‐up event in Ecuador comprising widespread ig‐
nimbrite eruption and batholith construction within an overall tensional regional stress field. Initiation of the regional flare‐up event coincides in time with a significant acceleration of Farallon/Nazca‐South Amer‐
ica convergence rates suggesting a positive feedback between faster plate convergence, asthenospheric melt production, mantle‐crust melt flux, and upper crustal arc magmatic productivity in Ecuador. Chapter 3 presents Re‐Os molybdenite and U‐Pb titanite ages related to mineralization and alteration fea‐
tures of latest Oligocene and Miocene porphyry‐related ore deposits in Ecuador. The new geochronologic data allow us to infer that the Miocene metallogenic belt of northern‐central Peru extends northwards into southern Ecuador, and potentially further north until Colombia. Miocene mineralization closely follows the distribution of Miocene arc magmatism in Ecuador. The regional spatio‐temporal distribution of porphyry Cu and associated epithermal mineralization within the Ecuadorian and Peruvian metallogenic belt seg‐
ments is similar. Intersections of Andean (NNE‐) trending structures with arc‐transverse faults and linea‐
ments related to suture zone geometries and block rotation in southern Ecuador represent highly prospec‐
tive sites for Miocene mineralization. The lack of Quaternary arc volcanic cover sequences and overall fa‐
vorable erosion levels are key parameters to preserve and expose widespread Miocene porphyry‐related mineralization in southern Ecuador. In the Western Cordillera of Ecuador porphyry‐related mineralization has locally been preserved, whereas the deeply eroded cores of porphyry systems are exposed at other lo‐
cations and significant parts of the mineralization have been removed. Porphyry systems are often spatially associated with intrusive clusters of batholith dimension, and formed towards the final stages of batholith assembly. Extensive shallow crustal magmatism during peaks of batholith construction may thus be disad‐
vantageous for the formation and preservation of porphyry‐related ore deposits, whereas favorable petro‐
genetic preconditioning of potential porphyry parental melts may occur towards the final phase of batho‐
lith construction. A direct spatio‐temporal association of Miocene porphyry‐related ore deposits and ridge (seamount chain) subduction or pulses of regional compression, as proposed elsewhere in the South American Andes, is not observed in Ecuador. The crustal basement of Ecuador comprises a collage of mostly Paleozoic‐Mesozoic tectono‐stratigraphic units of both continental and oceanic affinity in the Eastern Cordillera, and oceanic plateau units in the Western Cordillera and forearc region which were accreted in the Late Cretaceous. The diffuse paleo‐
v
continental suture zone is situated between the Eastern and Western Cordillera ranges where basement units are covered by Tertiary‐Quaternary arc volcanics. Chapter 4 presents a set of 58 new whole‐rock Sr, Nd, and Pb isotopic compositions of Late Oligocene and younger arc volcanics and associated intrusions of the Western Cordillera, its western foothills, and the central‐southern Ecuadorian Sierra region. Combining this new dataset with existing data on Quaternary arc volcanoes allows us to trace basement units of the Late Cretaceous suture zone at depth. Quaternary arc volcanics define distinct isotopic (Sr, Nd, Pb) groups for volcanoes situated east and west of the regional, roughly margin‐parallel Peltetec Fault, respectively. Late Oligocene to Late Miocene arc volcanics and intrusions of the southern Ecuador Sierra region overlap isotopically with recent arc volcanics in northern Ecuador suggesting along‐arc continuity of similar base‐
ment units at depth. Crustal isotopic contamination of Tertiary‐Quaternary arc magmas mainly takes place at deep to mid‐crustal levels except for granitoids of the Cangrejos‐Zaruma intrusive belt in southern Ecua‐
dor, where additional prominent shallow crustal assimilation is recorded. Isotopic compositions of arc magmas in northern‐central Ecuador follow a systematic across‐arc pattern where they evolve towards progressively more radiogenic 87Sr/86Sr and 207Pb/204Pb, and less radiogenic 143Nd/144Nd compositions at deep to mid‐crustal levels with increasing distance from the trench. This is consistent with regional, east‐
directed underthrusting of accreted oceanic plateau material along a broad suture zone below the paleo‐
continental margin as previously inferred from seismic studies. Chapter 5 presents a comprehensive dataset of the geochemical composition of Late Oligocene to Late Miocene intrusions and arc volcanics associated in space and time with porphyry‐related mineralization in Ecuador, focusing on the spatio‐temporal distribution pattern of adakite‐like geochemical features and ex‐
ploring their significance for porphyry‐related mineralization. The overall spatio‐temporal distribution of adakite‐like features in Ecuadorian arc magmas associated with porphyry systems is semi‐systematic; the relative proportion of adakite‐like (high Sr/Y) magmas increases with decreasing age, and is higher in northern‐central than in southern Ecuador. Broadly increasing Sr/Y and Sm/Dy ratios through time are con‐
sistent with progressively increasing high‐pressure crustal magma differentiation. High Sr/Y magmatism in the Late Tertiary is mainly due to strong Y (and heavy REE) depletion of parental melts at broadly constant Sr contents, where the former is related to fractionation/restite equilibration effects of amphibole, garnet, and titanite. While amphibole (± accessory titanite) fractionation/restite equilibration caused silicic melts to evolve towards adakite‐like compositions in the Early to Mid‐Miocene, combined amphibole and garnet fractionation/restite equilibration in Ecuadorian arc magmas has only been widespread since the Late Mio‐
cene and continues to the present day. A preferential association of adakite‐like features with a specific basement lithology cannot be observed. Increasing crustal thickness favorably influences the occurrence of adakite‐like features on a regional scale, but the latter are further modulated by a set of parameters which dynamically control mineral stabilities and mineral‐melt partitioning coefficients at a local scale. These include magma evolution depth (pressure) in a given crustal column and melt composition (degree of differentiation and melt water content). Por‐
phyry‐related deposits in Ecuador are often associated with intrusive clusters recording multi‐m.y. precur‐
sor magmatism where porphyry emplacement commonly represents a late intrusive event. Porphyry pa‐
rental melts tend to evolve towards more adakite‐like compositions than precursor intrusions if a signifi‐
cant relative age differences with respect to their emplacement exist. In contrast, systematic compositional changes between porphyry and precursor intrusions are not recorded if the relative age difference be‐
tween their respective emplacement events is small. As such, compositional changes between porphyry and precursor magmatism mostly reflect broad changes in arc magma composition through time at a re‐
gional scale. The fact that porphyry‐related ore deposits in Ecuador formed throughout the Late Oligocene to Late Miocene (24‐6 Ma) over a large latitudinal range (c. 0° to 3°30’S) supports the notion that any arc magma of a sufficient volume has the potential to form porphyry‐related mineralization. In some cases adakite‐like magmatism may, however, reflect favorable tectonomagmatic preconditioning of porphyry parental melts for subsequent porphyry‐related mineralization. vi Resumen Esta tesis de PhD presenta resultados geocronológicos en rocas del arco magmático del Terciario Tardío asociadas a depósitos de pórfidos y epitermales (Capítulo 2) y el tiempo de la mineralización y alteraciones asociadas a las intrusiones (Capítulo 3) así como datos isotópicos (Capítulo 4) y geoquímicos (Capítulo 5) de rocas volcánicas y plutónicas del arco Oligoceno‐Mioceno asociadas con sistemas porfiríticos. Resúmenes específicos de cada capítulo preceden cada parte de la tesis. Una introducción general (Capítulo 1) explica algunas de las ideas generales del proyecto y las conclusiones generales se encuentran en el Capítulo 6. Ecuador presenta un gran número de granitoides y rocas volcánicas relacionadas con el arco magmático Terciario cuyas características geocronológicos han sido mayormente basadas en edades K‐/Ar y en trazas de fisión en zircón. En el Capítulo 2 se presenta el primer set de datos U/Pb en zircón realizado a escala re‐
gional, obtenidos en múltiples plutones e intrusiones porfiríticas en la Cordillera Occidental del Ecuador y en la Sierra Austral, lo que nos ha permitido una calibración robusta de la historia intrusiva durante el Ter‐
ciario en Ecuador. Las dataciones en zircón fueron obtenidas a partir de análisis de espectrometría de masa a partir de disolución isotópica ‐ ionización termal y también a partir de ablación laser. Excepto por una muestra, todas las demás analizadas carecen completamente de zircones heredados lo que sugiere un ba‐
samento oceánico pobre en zircones; la presencia de antecristales es relativamente abundante en muchos plutones lo que indica que granitoides mas antiguos pudieron haber sido reciclados. En rocas plutónicas donde se han obtenido edades K/Ar y U‐Pb se puede observar que los resultados son concordantes entre 1‐4 Ma, lo que sugiere que las edades K/Ar pueden ser usadas como un proxi semi‐
exacto para las rocas del arco magmático en Ecuador a escala regional. Distribuciones espacio temporales de la rocas magmáticas terciarias indican un evento de "flare‐up" desde el Oligoceno Tardío al Mioceno Temprano/Medio el que consistió en una importante y amplia erupción de ignimbrita y así mismo una in‐
tensa formación de batolitos durante un periodo de distensión regional. La iniciación del evento "flare‐up" coincide en el tiempo con una importante aceleración de la convergencia entre las placas Farallón/Nazca – Sudamérica sugiriendo que hay una relación directa entre la convergencia rápida de placas tectónicas, la producción de fundidos astenosféricos, flujo de fundidos manto‐corteza y actividad magmática en la corte‐
za superior del arco. En el Capítulo 3 se presenta edades Re‐Os en molibdenita y U‐Pb en titanita las cuales están relacionadas a mineralizaciones y alteraciones de los depósitos de tipo pórfido y epitermales del Oligoceno Tardío y del Mioceno. Estos nuevos resultados geocronológicos nos permiten inferir que el cinturón metalogénico de los Andes Nor‐centrales del Perú se extendieron al norte hacia el Sur de Ecuador e incluso posiblemente hasta Colombia. La mineralización del Mioceno siguen un patrón de distribución muy parecido a la distribu‐
ción del arco magmático Micénico en Ecuador. La distribución espacio‐temporal de los pórfidos de Cu y mi‐
neralizaciones epitermales asociadas dentro de los cinturones metalogénicos del Ecuador y Perú es muy similar. Las zonas donde se intersecan estructuras de rumbo Andino (NEE) con fallas transversales al arco y lineaciones relacionadas con zonas de sutura y/o bloques rotados en el Sur de Ecuador, representan zonas altamente prospectivas para mineralizaciones miocénicas. La ausencia de cobertura volcánicas cuaternaria y sobre todo la actividad erosiva son factores claves para la preservación y la exposición de las rocas miocénicas de tipo‐pórfido en el Sur de Ecuador. En la Cordillera Occidental del Ecuador las mineralizaciones asociadas al pórfido han sido localmente preservadas, mientras en unas zonas se puede observar el núcleo del sistema de pórfido, en otras zonas las partes más importan‐
tes del sistema han sido removidas debido a una mayor erosión. Los sistemas de pórfidos están a menudo asociados con intrusivos dispersos que se formaron en las etapas finales de la formación de grandes batoli‐
tos. Un magmatismo intenso en la corteza superior durante los mayores picos construcción de batolitos posiblemente no es favorable para la formación y preservación de los depósitos asociados con pórfidos, mas bien las mejores condiciones petrogenéticas se dan cuando los fundidos que forman los pórfidos pue‐
vii
den ocurrir hacia la parte final de la formación de los batolitos. En Ecuador no se observa una relación es‐
pacio‐temporal directa entre la edad de los depósitos tipo pórfido y la subducción de ridges oceánicos y/o pulso de compresión regional, como se ha propuesto en muchas otras partes en Sudamérica. El basamento cortical en Ecuador consiste de una serie de unidades tectono‐estratigráficas del Paleozoico‐
Mesozoico de afinidad continental y oceánica en la Cordillera Oriental y rocas relacionadas con un plateau oceánico en la Cordillera Occidental y la costa del Ecuador las cuales fueron acrecionadas en el Cretácico Tardío. La sutura paleocontinental es muy difusa y estaría situada entre la Cordillera Oriental y Occidental donde las unidades del basamento están cubiertas por rocas volcánicas del Arco Terciario‐Cuaternario. En el Capitulo 4 se presenta un set de 58 datos isotópicos de Sr, Nd, Pb en roca total llevadas a cabo en rocas volcánicas e intrusivas del Oligoceno Tardío y más jóvenes, expuestas en la Cordillera Occidental y en la Sie‐
rra Central y Austral del Ecuador. Al combinar estos datos isotópicos con datos existentes en rocas volcáni‐
cas del Cuaternario nos da indicios del basamento que nos permiten trazar posibles zonas de sutura en la profundidad. Rocas volcánicas Cuaternarias definen distintos grupos isotópicos (Sr, Nd, Pb) para volcanes que están si‐
tuados ya sea al este y al oeste de la falla regional de Peltetec. Rocas volcánicas e intrusivas del Oligoceno Tardío al Mioceno tardío en la Sierra Sur del Ecuador se sobrelapan isotópicamente con rocas volcánicas del arco actual presente solamente en el Norte de Ecuador, sugiriendo esto que existe una continuidad en las unidades profundas del basamento a lo largo del arco. Contaminación isotópica cortical de rocas del arco Terciario‐Cuaternario toma lugar a niveles de la corteza media‐profunda, excepto para los granitoides del cinturón intrusivo de Cangrejos‐Zaruma donde se registra una prominente contaminación cortical de corte‐
za superior. Composiciones isotópicas del arco magmático en la Sierra Norte y Centro siguen una sistemáti‐
ca a través del arco donde estas evolucionan progresivamente hacia valores 87Sr/86Sr y 207Pb/204Pb mas ra‐
diogénicos y composiciones de 143Nd/144Nd menos radiogénicas a profundidad de corteza media conforme se incrementa la distancia desde la fosa. Esta observación es consistente con la presencia de rocas del pla‐
teau oceánico a lo largo de una amplia zona de sutura bajo el margen paleocontinental como ha sido inferi‐
do ya por estudios sísmicos. En Capítulo 5 se presenta un set de datos completo y exhaustivo consistente en composiciones geoquími‐
cas de las intrusiones y rocas volcánicas de arco del Oligoceno Tardío‐Mioceno Tardío, las cuales están aso‐
ciados en el espacio‐tiempo con mineralizaciones tipo pórfido en Ecuador, el cual se enfoca en el patrón de distribución espacio‐temporal de características geoquímicas adakíticas y explora además la importancia que estas tienen para las mineralizaciones tipo pórfido. La distribución total espacio‐temporal de las carac‐
terísticas adakíticas en las rocas magmáticas de arco que tienen sistemas porfiríticos asociados es semi‐
sistemática; la proporción relativa de magmas adakíticos (valores de Sr/Y elevados) se incrementa cuando la edad decrece y es mayor en la región Norte y Centro en comparación con la zona Sur. Un amplio incre‐
mento en los valores Sr/Y y Sm/Dy a través del tiempo es consistente con un incremento progresivo de la diferenciación magmática a niveles corticales de presión elevada. Valores altos de Sr/Y en rocas magmáticas del Terciario Tardío es principalmente debido a un fuerte empo‐
brecimiento de Y (y de HREE) en los fundidos parentales cuando se tiene concentraciones aproximadamen‐
te constantes de Sr, en las cuales el Y está relacionado con los efectos de fraccionamiento (y/o equilibrio en la restita) de anfíbol, granate y titanita. Mientras el fraccionamiento (y/o equilibrio en la restita) del anfíbol (+/‐ titanita como accesorio) causa que los fundidos silícicos evolucionen hacia composiciones tipo adakíti‐
cas en el Mioceno Temprano‐Medio, el fraccionamiento (y/o equilibrio en la restita) del anfíbol y el granate combinados, ha sido solo ampliamente existente desde el Mioceno Tardío y continua hasta la actualidad. No se observa una asociación preferencial de las características adakíticas con un tipo específico de litología del basamento. Un incremento del grosor de la corteza influencia favorablemente la ocurrencia de carac‐
terísticas adakíticas a escala regional, pero esta última es posteriormente modulada por una serie de pará‐
metros los cuales controlan dinámicamente las estabilidades minerales y los coeficientes de partición mine‐
ral‐fundido a una escala local. Estos parámetros incluyen la profundidad de evolución del magma (presión) viii en una columna cortical dada y además incluye la composición del fundido (grado de diferenciación y con‐
tenido de agua del fundido). Depósitos de tipo pórfido y epitermales en Ecuador están a menudo asociados con intrusivos dispersos los cuales graban un magmatismo precursor (a la escala de varios millones de años) donde el emplazamiento del pórfido normalmente representa el último evento de intrusión. Los fundidos parentales que forman los pórfidos tienden a evolucionar hacia composiciones mas adakíticas que las intru‐
siones precursoras si hay una diferencia significativa en las edades de su emplazamiento. Al contrario, no hay evidencia de un cambio composicional entre el pórfido y sus intrusiones precursoras si es que la edad relativa entre sus respectivos emplazamientos es pequeña. De esta manera, los cambios composicionales entre el pórfido y el magmatismo precursor, refleja mayormente cambios grandes a lo lar‐
go del tiempo en la composición magmática del arco. El hecho de que los depósitos de tipo pórfido y epi‐
termales en Ecuador se formaron durante el Oligoceno Tardío‐Mioceno Tardío (24‐6 Ma) a lo largo de un gran rango latitudinal (c. 0° a 3°30’S), soporta la idea de que cualquier magma de arco con un volumen sufi‐
ciente tiene el potencial de formar mineralizaciones económicas. En algunos casos los magmas adakíticos pueden sin embargo, reflejar pre‐condiciones tectono‐magmáticas favorables de los fundidos parentales del pórfido para una subsecuente mineralización relacionada con el pórfido. ix
x Résumé étendu Le travail de cette thèse présente des données géochronologiques concernant la mise en place des intrusions d’âge Tertiaire supérieures en Equateur, et sur les événements de minéralisa‐
tion et d'altération hydrothermale liés à ses in‐
trusions, ainsi que la composition isotopique et géochimique des roches volcaniques et plutoni‐
ques associées à l’arc d’âge Oligocène‐Miocène. Les granitoïdes d'âge Oligocène‐
Miocène et l'influence des facteurs géodynamiques pour l'arc magmati‐
que en Equateur L’Equateur hôte un très grand nombre de grani‐
toïdes et des roches volcaniques liées à l’arc d’âge Tertiaire, dont les caractéristiques géo‐
chronologiques sont principalement basées sur les données de la méthode de datation K‐Ar (et les traces de fissions sur des zircons). Le chapitre 1 de cette thèse présente les premières données régionales des âges U‐Pb des zircons provenant des intrusions plutoniques et porphyriques de la cordillère Ouest et le sud de la région de Sierra (Fig. 1), permettant une calibration temporale robuste de l’histoire intrusive d'âge Tertiaire en Equateur. Les zircons étaient datés avec deux méthodes, soit par TIMS (spectrométrie de masse par ther‐
mo‐ionisation), soit par LA‐MC‐ICP‐MS (émission en plasma induit couplée à la spectrométrie de masse et laser ablation). Du nord au sud, les âges suivants ont été obtenus: 12.87±0.05 Ma pour la batholite d’Apuela à Cuellaje, pénétrée par des dykes porphyriques de 9.01±0.06 Ma à Junin; 25.5±0.7 Ma pour le pluton de Telimbela centra‐
le; 21.46±0.09 Ma et 21.22±0.17 Ma pour le plu‐
ton Balsapamba et une intrusion du dyke porphy‐
rique dans la zone d’El Torneado respectivement; 14.84‐15.33 Ma pour le batholite de Chaucha et 9.79±0.03 Ma pour une intrusion porphyrique à Tunas; 20.26±0.07 Ma et 19.89±0.07 Ma pour les deux porphyres à Gaby‐Papa Grande; 7.13±0.07 Ma pour un dôme intra‐caldeira du centre volca‐
nique de Quimsacocha; 30.7±0.7 Ma pour une roche subvolcanique de Saraguro à Tres Chorre‐
ras; 16.04±0.04 Ma pour une intrusion porphyri‐
que à El Mozo ; 26.0±0.7 Ma pour un pluton à Cangrejos; 20.7±0.9 Ma pour un pluton au nord d’Zaruma; 24.04±0.07 Ma pour une intrusion porphyrique à Portovelo, et 92.0±1.6 Ma pour une intrusion porphyrique à Curiplaya. Sauf l'échantillon de Tres Chorreras, dans toutes les intrusions (ou les roches sub‐volcaniques) étudiées, les zircons hérités sont complètement absents suggérant la base de la croûte est océa‐
nique et appauvrie en zircons; les zircons de type "antecryst" sont relativement abondants dans plusieurs plutons indiquant que le recyclage d’un ancien arc avait pris place. Lorsque les données des deux méthodes de datation géochronologi‐
ques, K‐Ar et U‐Pb, coexistent pour un système intrusif, les âges obtenus par les différentes mé‐
thodes sont généralement concordant et dévoi‐
lent une différence d'âge entre 1‐4 Ma au maxi‐
mum, ce qui implique que les âges de K‐Ar peu‐
vent être utilisé comme des données semi‐
précises pour l’arc magmatique d’âge Tertiaire en Équateur à l'échelle régionale. La répartition spatio‐temporelle des intrusions et roches volcaniques d'âge Tertiaires indique un vif échauffement de magmatisme d’âge Oligocène terminal jusqu’au Miocène inférieure et moyen en Équateur, comprenant l'éruption régionale des ignimbrites et la construction des batholites dans un champ de contrainte de tension, généra‐
lement à une pente constante du slab. L’initiation de l’événement d’échauffement régionale coïnci‐
de dans le temps avec une accélération impor‐
tante dans le taux de convergence des plaques Sud Amérique et Farallon/Nazca, suggérant une relation positive entre une convergence très ra‐
pide des plaques, génération de la fusion partielle dans l'asthénosphère, le flux de la fusion entre manteau et croûte, et la productivité de l'arc magmatique en Équateur. Cela pourrait être envisagée par un taux plus éle‐
vé du fluide généré du slab subducté dans un vo‐
lume donné dans le coin mantellique, et/ou par un changement dans la dynamique des flux as‐
thénosphériques où des anomalies thermiques positives se développent dans le coin mantellique et les taux de reconstitution du matériel du man‐
teau fertile augmente en réponse du retour xi
Figure 1: Carte géologique de la région des Cordillères d'Equateur, montrant les éléments géologiques principaux de l'Equateur et l'arc magmatique d'âge Tertiaire. La figure au‐dessous à gauche montre la situation géodynamique du bassin de Panama. Adapté par Litherland et al. (1994), Steinmann (1997), Dunkley & Gaibor (1997), McCourt et al.
(1997), Pratt et al. (1997), Hughes et al. (1998), Meschede & Barckhausen (2001), et Palacios et al. (2008). xii du flux induit. L'intensification des transferts de chaleur de la conduction et l’advection dans la croûte pourrait déclencher un processus de rela‐
tion positive tectono‐magmatique et thermique, qui peut faciliter l’intensification de la fusion par‐
tielle de la croûte et le volumineux stockage du magma à des niveaux supérieurs, menant à la construction des batholites, et/ou l’éruption des ignimbrites en Équateur au cours de l'Oligocène terminal à Miocène inferieur‐moyen. La ceinture métallogénique d'âge Miocène en Equateur Le chapitre 2 présente les âges calculés grâce aux méthodes géochronologiques dans la molybdéni‐
te selon le système Re‐Os et dans la titanite selon le système U‐Pb; ces âges sont liés à la minérali‐
sation/altération hydrothermale associée avec des intrusions des gisements de porphyre cupri‐
fère ou épithermaux d’âge Oligocène et Miocène en Équateur. Figure 2: Distribution spatio‐temporelle des gisements d'âge Miocène associés aux intrusions d'âge Oligocène‐Miocène en Equateur. Il y a de pics de la minéralisation dans le Miocène inférieur avec un deuxième pic dans le Miocène supé‐
rieur. Dans un système batholitique donné (Apuela, Balsapamba‐Telimbela, Chaucha) la minéralisation se manifeste dans un stage final de l'évolution magmatique du batholite par rapport à son construction initiale (après c. 5‐15 m.y.). Pour comparaison les pulses compressives (I = Inca; Q = Quechua) en Equateur (boxes noires; Hungerbühler et al. 2002) et en Pérou (boxes grises: Noble & McKee 1999; boxes blanches: Benavides‐Cáceres 1999) sont montrées à droite.
xiii
La molybdénite associée à l’altération potassique et phylliteuse à Junin (gisement porphyre Cu‐Mo) a donné des âges de 6.63±0.04 Ma et 6.13±0.03 Ma. Les âges calculés selon le système chronolo‐
gique Re‐Os dans les molybdénites associées à l’altération potassique dans les systèmes porphy‐
riques de Telimbela et Balsapamba sont de 19.2±0.1 Ma et 21.5±0.1 Ma, respectivement. Dans les systèmes porphyriques de Cu‐Mo à Chaucha, les âges obtenus dans des molybdénites associées aux altérations potassiques et phylli‐
teuses selon le système Re‐Os sont de 9.92±0.05 Ma (à Tunas‐Naranjos) et 9.5±0.2 Ma (à Gur‐Gur), respectivement. Dans le système porphyrique Au‐Cu de Gaby, un âge de 20.6±0.1 a été calculé pour la molybdénite selon le système chronologi‐
que Re‐Os pour des brèches hydrothermales sul‐
furées (éventuellement associés à une altération phylliteuse), et un âge de 20.17±0.16 Ma selon le système chronologique U‐Pb pour une titanite associé à l’altération de type Na‐Ca a été obte‐
nue. Au gisement polymétalliques de Tres Chor‐
reras, les âges obtenus selon le système chrono‐
logique Re‐Os dans les molybdénites sont à 12.93±0.07 Ma et 12.75±0.07 Ma, et sont asso‐
ciées à une brèche hydrothermale liée à une in‐
trusion et une veine polymétalliques, respecti‐
vement. La molybdénite associée à l'altération de type Na‐Ca dans le système porphyrique Au‐Cu de Cangrejos a donné un âge de 23.5±0.1 Ma. Les nouvelles données géochronologiques nous permettent de déduire que la ceinture métallo‐
génique d’âge Miocène au nord du Pérou s'étend du nord vers le sud de l'Équateur, et potentielle‐
ment plus au nord jusqu'à la Colombie. La miné‐
ralisation d’âge Miocène suit de près la réparti‐
tion de l’arc magmatique d’âge Miocène en Equa‐
teur, qui se caractérise souvent par des déforma‐
tions syn‐magmatiques. Sur une échelle régiona‐
le, la distribution spatio‐temporelle des gise‐
ments de porphyre cuprifères et épithermaux associée dans les segments de la ceinture métal‐
logénique équatorienne et péruvienne est simi‐
laire. L’intersection des structures Andines (NNE‐) avec de failles transversales et des traits liés à des géométries de zones de sutures et de rotation de blocs dans le sud de l'Equateur représentent des xiv sites très favorables pour la minéralisation. L'ab‐
sence de la couverture volcanique quaternaires et l'érosion globale des niveaux favorables sont des paramètres clés pour préserver et exposer généralement les gisements de porphyres cupri‐
fères et épithermaux d’âge Miocène dans le sud de l'Équateur. Dans la Cordillère de l'Ouest de l'Équateur la minéralisation des porphyres cupri‐
fères à été localement préservée, tandis que les cœurs profonds érodés des systèmes porphyri‐
ques sont exposés à d'autres endroits et des par‐
ties importantes de la minéralisation ont été en‐
levés. Alors que les complexes de batholites peuvent marquer structurellement les sites favorables à la minéralisation, de vaste magmatisme dans la croûte supérieure pendant les pics de la cons‐
truction des batholites peut être désavantageux pour la formation et la préservation des gise‐
ments de porphyre cuprifères. En revanche, le préconditionnement pétrogénétique favorable des liquides porphyriques parentaux peut se pro‐
duire vers la phase finale d'assemblage des ba‐
tholites. Mais peut‐être applicable pour quelques gisements, une générale association spatio‐
temporelle entre la formation de gisements et les pulses de compression régional ou la subduction des chaînes montagneuses sous‐marines n'est pas observée en l'Équateur (Fig. 2). La composition isotopique des mag‐
mas d'âge Oligocène‐Miocène et les domaines isotopiques de la croûte équatorienne En Equateur, dans la Cordillère Orientale, la base de la croûte comprend un ensemble d’unités tec‐
tono‐stratigraphiques de la plupart d’âge Paléo‐
zoïque et Mésozoïque d'affinité à la fois conti‐
nentale et océanique. Des unités du plateau océanique accumulés au Crétacé supérieur for‐
ment la base de la croûte dans la Cordillère Occi‐
dentale et la région d’avant‐arc. La zone de sutu‐
re paléo‐continentale est située entre l'Est et l'Ouest de la série de Cordillère où des unités de la base de la croûte sont couverts par des roches volcaniques du Tertiaire et Quaternaire. Un grand nombre d'information sur la géochimie des ro‐
ches volcaniques du Quaternaire existe pour l'arc de la Zone Volcanique du Nord dans le nord de l'Equateur, alors que les séquences de la couver‐
ture Tertiaire en Équateur centrale et sud sont mal caractérisées isotopiquement. Le chapitre 3 présente un ensemble de 58 nouveaux résultats de roche totales des compositions isotopiques de Sr, Nd et Pb des échantillons Tertiaire de l’arc volcanique et des intrusions associées de la Cor‐
dillère Occidentale et du sud de la région de la Sierra équatorienne. La combinaison de ce nouvel ensemble de données avec les données existan‐
tes sur l’arc volcanique Quaternaire nous permet de suivre les unités basales de la croûte de la zo‐
ne de suture du Crétacé terminal en profondeur. L’arc volcanique d’âge Quaternaire définit des groupes isotopiques distincts pour des volcans situés à l'est et à l'ouest de la faille Peltetec. Les roches volcaniques et les intrusions de la région sud équatorienne de Sierra d’âge Oligocène ter‐
minal‐Miocène terminal chevauchent isotopi‐
quement avec des roches volcaniques récentes à l’est de la faille Peltetec dans le nord de l'Equa‐
teur suggérant ainsi une continuité des unités basales de la croûte similaires en profondeur. Les granitoïdes d’âge Oligocène‐Miocène de la Cordillère Occidentale et sur ses contreforts occi‐
dentaux montrent des compositions isotopiques les plus primitives de Sr et Nd identifiées à ce jour dans l’arc magmatique équatorien d’âge Tertiai‐
re‐Quaternaire; les unités primitives du plateau océanique du Crétacé constituent leurs assimilant en profondeur, entraînant de ces magmas d’arc de devenir plus primitifs isotopiquement avec l'assimilation de la croûte. La contamination isotopique crustale du magma d’arc d’âge Tertiaire‐Quaternaire a lieu principa‐
lement à des niveaux profonds à moyens de la croûte à l'exception de la ceinture intrusive des granitoïdes du Cangrejos‐Zaruma dans le sud de l'Équateur, où plus éminente assimilation de ma‐
tériau de la croûte supérieure est caractérisé par une composition de Sr et Pb très radiogéniques, et une composition isotopique de Nd peu radio‐
géniques (Fig. 3). Les compositions isotopiques du Sr, Nd et Pb des magmas d'arc dans le nord‐
central de l'Équateur suivent un schéma systéma‐
tique dans l'ensemble d’arc où ils évoluent pro‐
87
gressivement vers Sr/86Sr et Figure 3: Diagrammes de 87Sr/86Sr, εNdinitial, 206Pb/204Pb, et 207Pb/204Pb vs. Sr/Y; Sr/Y est utilisé comme indica‐
teur d'évolution magmatique dans le croûte supérieur vs. inférieur (Sr/Y >30 pour le dernier). Dans l'ordre de leurs distribution géographique, les centres magmati‐
ques forment des groupes isotopiques subparallèles pour des valeurs de Sr/Y >30 ce qui implique que les processus de AFC (assimilation et cristallisation frac‐
tionnelle) incluent des unités différents à la base de la croûte. En plus, il y a d'évolution magmatique dans la
croûte supérieur pour le Tertiaire supérieur (particuliè‐
rement pour les intrusions de Cangrejos‐Zaruma) dont l'assimilation comprend des lithologies caractérisées par des compositions plus radiogéniques en Sr et moins radiogéniques en Nd. xv
207
Pb/204Pb plus radiogéniques, et une composi‐
tion moins radiogénique en 143Nd/144Nd du pro‐
fond au moyen niveau de la croûte avec la crois‐
sance de la distance depuis la tranchée. Ceci est cohérent avec la poussée régionale du matériel des plateaux océanique accrété le long d'une vas‐
te zone de suture en dessous de la marge paléo‐
continentale comme précédemment déduit à partir des études sismiques. La géochimie des granitoïdes et ro‐
ches volcaniques d'âge Oligocène‐
Miocène et la distribution temporelle des compositions "adakite‐like" Enfin, le chapitre 4 présente un ensemble de données de l’Oligocène terminal au Miocène su‐
périeur de la composition géochimique des intru‐
sions associées dans l'espace et dans le temps avec les porphyres cuprifères et la minéralisation épithermal en Equateur, complétées par des données en plusieurs formations de l'arc volcani‐
que du même âge. Notre objectif est de décrire la distribution spatio‐temporelle des compositions "adakite‐like" dans le contexte de l'évolution géochimique d'arc du Tertiaire terminal, et d'ex‐
plorer son importance pour les intrusions asso‐
ciées à la minéralisation cuprifère/aurifère en Équateur. La plupart des intrusions représentent des tonali‐
tes, granodiorites et diorites quartzeuse modé‐
rément à fortement différenciées et portant de l'hornblende ± biotite, et font souvent partie de plus grands complexes batholitiques d’âge Oligo‐
cène‐Miocène; les roches volcaniques allant
Figure 4: Diagrammes des éléments de traces et leurs rapports vs. l'âge. La distribution des éléments de traces impli‐
que un épaississement progressif de la croûte pendant l'Oligocène‐Miocène. En plus, des facteurs pétrogénétiques (fractionnement de l'amphibole) contrôlent l'appauvrissement extrême en Y pour les compositions magmatiques silici‐
ques. xvi d'une composition andésitique à dacitique‐
rhyolitique. L'ensemble spatio‐temporel de la distribution des magmas de type "adakite‐like" dans l'arc équatorien est semi‐systématique; la proportion relative des magmas "adakite‐like" augmente avec la diminution de l'âge, et est plus élevé dans le centre‐nord que dans le sud de l'Équateur. Les centres magmatiques caractérisée par, en partie, un magmatisme "adakite‐like" sont principalement encaissés par la Cordillère Occidentale et comprennent Balsapamba (c. 21 Ma), Apuela‐Junin (13‐6 Ma), Chaucha (vers 10 Ma), et Quimsacocha (7 Ma). Les caractéristiques "adakite‐like" (haute Sr/Y) des magmas d’arcs équatoriens d’âge Tertiaire terminal sont princi‐
palement dues à un fort appauvrissement en Y (et les terres rares lourdes) de leurs liquides pa‐
rentaux en gardant les teneurs plus ou moins constantes en Sr, et sont liés au fractionne‐
ment/effets d’équilibration de restite d'amphibo‐
le, grenat, et titanite. Du Miocène inférieur au moyen, le fractionne‐
ment/l'équilibration de restite d’amphibole (± titanite comme accessoire) a causé l’évolution du liquide silicique vers une composition "adakite‐
like". La combinaison du fractionnement/ l’équilibration de restite d’amphibole et grenat dans les magmas d'arc équatorien n'a été généra‐
lisée que depuis le Miocène supérieur. L'appau‐
vrissement de Y par fractionnement/équilibration de restite d'amphibole est très efficace mais seu‐
lement pour les liquides de compositions silici‐
ques. Par contre, le fractionnement/équilibration de restite de grenat produit un appauvrissement fort de Y déjà dans les liquides de compositions plus mafiques, c'est à dire, pendant la phase ini‐
tiale de différenciation. Le fractionnement des plagioclases dans la croûte supérieure affecte certains, mais pas tous les magmas d’arc d’âge Tertiaire dans le sud de l'Equateur; il est d'une importance pétrogénétique mineure pour les in‐
trusions d’âge Miocène de la Cordillère Occiden‐
tale dans le nord‐centre de l'Equateur. Une asso‐
ciation de caractéristiques préférentielles des
Figure 5: Distribution de Sr/Y vs. Y pour les intrusions
porphyriques d'âge Tertiaire tardif et les roches intru‐
sives phanéritiques associées avec les porphyres
en Equateur. Sauf le porphyre de Cangrejos, tous les porphyres ne montrent pas d'évidence pour l'évolution des liqui‐
des parentaux dans la croûte supérieure (fractionnement de plagioclase). Par contre, les liquides parentaux (sauf Can‐
grejos et Gaby) sont caractérisé par le fractionnement de l'amphibole ± titanite ± grenat aux niveaux plus profonds de la croûte (et/ou les liquides parentaux sont plus riches en H2O) en montrant des compositions "adakite‐like". xvii
compositions "adakite‐like" avec une lithologie de la base de la croûte ne peut être observée. Une variation systématique des éléments traces (Sr, Y, REE) dans le temps sont révélateurs d’un progressif épaississement de la croûte équato‐
rienne de l'Oligocène terminal au Miocène termi‐
nal (Fig. 4). Tout en augmentant l'épaisseur de la croûte influence favorablement l'apparition des compositions "adakite‐like" à l'échelle régionale; ces derniers sont en outre modulées par un en‐
semble de paramètres qui contrôlent dynami‐
quement les stabilités des minéraux et les coeffi‐
cients de partage minéral‐liquide à l'échelle plu‐
tôt locale. Il s'agit notamment de l'évolution du magma en profondeur (pression) dans une co‐
lonne donnée de la croûte et la composition du liquide (degré de différenciation et la teneur en eau). Celui‐ci indique notamment la migration du magmatisme crustal à une plus grande profon‐
deur, et/ou en augmentant le contenu en eau du système magmatique. Des changements systématiques dans la compo‐
sition entre les intrusions porphyriques et des précurseurs intrusifs ne sont pas enregistrés si la différence de temps entre leur mise en place est faible (Fig. 5). Le fait que la minéralisation de type porphyre cuprifère et épithermal "high sulfida‐
tion" en Équateur existe de l'Oligocène terminal entier à Miocène supérieur (24‐6 Ma) et une lar‐
ge latitude de grande taille (c. 0° à 3° 30'S) sou‐
tient l'idée que toutes les magmas d'un arc de volume suffisant ont le potentiel pour causer de minéralisation liée aux intrusions. Le magmatis‐
me de composition "adakite‐like" peut, cepen‐
dant, indiquer un favorable environnement tec‐
tono‐magmatique et un préconditionnement fa‐
vorable des liquides porphyriques parentaux pour la minéralisation de type porphyre cuprifère. Références Benavides‐Cáceres, V. (1999): Orogenic evolution of the Peruvian Andes: The Andean cycle. In: Skinner, B. J. (ed.), Geology and ore deposits of the central Andes. SEG Special Publication 7; 61–107. Dunkley, P. N. & Gaibor, A. (1997): Mapa geologico de la Cordillera Occidental del Ecuador entre 2°‐3° S. es‐
cale 1/200.000. CODIGEM‐Min. Energ. Min.‐BGS publs., Quito. xviii Hughes, R.A., Bermudez, R., Espinel, G. (1998): Mapa geológico de la Cordillera Occidental del Ecuador entre 0°‐1°S, escala 1:200.000. CODIGEM‐Ministerio de En‐
ergía y Minas‐BGS publs., Quito. Hungerbühler, D., Steinmann, M., Winkler, W., Sew‐
ard, D., Egüez, A., Peterson, D. E., Helg, U., Hammer, C. (2002): Neogene stratigraphy and Andean geodynam‐
ics of southern Ecuador. Earth Science Reviews 57; 75–
124. Litherland, M., Aspden, J. A., Jemielita, R. A. (1994): The metamorphic belts of Ecuador. Overseas Memoir 11. BGS, Keyworth. McCourt, W. J., Duque, P., Pilatasig, L. F. and Villago‐
mez, R. (1997): Mapa geológico de la Cordillera Occi‐
dental del Ecuador entre 1° ‐ 2° S., escala 1/200.000. CODIGEM‐Min. Energ. Min.‐BGS publs., Quito. Meschede, M. & Barckhausen, U. (2001): The relation‐
ship of the Cocos and Carnegie ridges: age constraints from paleogeographic reconstructions. International Journal of Earth Sciences 90; 386‐392. Noble, D. C. & McKee, E. H. (1999): The Miocene met‐
allogenic belt of central and northern Perú. SEG Spe‐
cial Publication 7; 155‐193. Palacios, O., Pilatasig, L., Sanchez, J., Gordon, D., Shaw, R., Feininger, T. (2008): Mapa geologico binacional region sur del Ecuador y norte del Peru. Ingeomin, 1 : 500,000. Pratt, W. T., Figueroa, J. F., Flores, B. G. (1997): Mapa geologico de la Cordillera Occidental del Ecuador entre 3°‐4°S. escale 1/200.000. CODIGEM‐Min. Energ. Min.‐
BGS publs., Quito. Steinmann, M. (1997): The Cuenca basin of southern Ecuador:tectono‐sedimentary history and the Tertiary Andean evolution. PhD Thesis, Institute of Geology ETH Zürich, Switzerland, 176 p. CHAPTER I THEORETICAL BACKGROUND AND AIMS OF THESIS Introduction Porphyry Cu deposits and associated epithermal and polymetallic vein mineralization, hereafter referred to as "porphyry‐related ore deposits", form as parts of hydrothermal systems related to fluid exsolution from shallow crustal intrusions (e.g., Seedorff et al. 2005). They can be found in subduction zone settings worldwide, and are re‐
garded as the product of a series of common‐
place geologic processes which have to combine favorably to eventually result in economic miner‐
alization (e.g., Tosdal & Richards 2001). Parental melts to porphyry intrusions ultimately derive from the supra‐slab mantle wedge, the latter fluxed by a volatile‐rich slab component. Subse‐
quent lower to mid‐crustal magma evolution of porphyry parental melts produces volatile‐
enriched andesitic or more differentiated melt compositions which then ascend to upper crustal levels where porphyry intrusive and associated hydrothermal systems may form. The localization of the latter is often structurally controlled (e.g., Richards 2003), although this is not always the case (Sillitoe 2000). As far as magma sources and early differentiation stages are concerned, a high fO2 is often inferred for porphyry parental melts with the potential to form high‐tonnage ore deposits (e.g., Sillitoe 2000; Mungall 2002). Depending on the prevail‐
ing fO2, S as a component of a silicate melt com‐
monly occurs in variable proportions of sulfide and sulfate where the melt S solubility is signifi‐
cantly higher if sulfate, rather than sulfide, forms the dominant S species (e.g. Jugo et al. 2005). If the S content of a silicate melt at supra‐liquidus conditions exceeds the melt S solubility, S‐rich immiscible liquids will form, which scavenge lithophile and chalcophile elements from the sili‐
cate melt (Jugo et al. 2005). If, by virtue of their higher density, S‐rich immiscible melt globules coalesce and fractionate from a given silicate melt batch, the silicate melt will become de‐
pleted in chalcophile metals (such as Cu and Au), thus decreasing its potential for subsequent por‐
phyry Cu‐Au mineralization. Although mass bal‐
ance calculations indicate that, in principle, any arc magma (average andesite Cu content 60 ppm; Cline & Bodnar 1991) may contribute to the for‐
mation of a porphyry‐related ore deposit as long as the integrated magmatic system comprises a sufficient volume (≥100 km3; Cline & Bodnar 1991), parental melts related to large porphyry‐
related ore deposits may be expected to be highly oxidized, sulfate‐dominated silicate melts whose metal budget has not been substantially depleted by metal loss to immiscible sulfide melts in or close to the magma source region (Sillitoe 2000; Mungall 2002). Porphyry‐related ore deposit distribution and size is diachronous through time and space, implying that certain tectonomagmatic environments are particularly prolific (or unfavorable) for maximiz‐
ing the potential of intrusion‐related mineraliza‐
tion. Spatio‐temporal clusters of porphyry‐
related ore deposits may be due to favorable ex‐
posure and preservation conditions, and/or in‐
creased rates of deposit formation (Wilkinson & Kesler 2009). Considering the latter, the in‐
creased abundance of porphyry‐related ore de‐
posits (or their overall larger tonnage) may reflect favorable tectonomagmatic settings which, in part, might derive from special geodynamic envi‐
ronments (Tosdal & Richards 2001). Amongst others, the subduction of seamount chains has been proposed to show a positive spatio‐
temporal correlation with porphyry‐style miner‐
alization (e.g., in the southern and central Andes; Rosenbaum et al. 2005; see also Cooke et al. 2005). Constraining potentially favorable condi‐
tions for porphyry‐related mineralization is of major interest for the design of regional mineral exploration campaigns for this type of ore de‐
posit. 1
In this context, Thieblémont et al. (1997) note that intrusions associated with porphyry‐style mineralization often tend to be of "adakitic" composition, potentially implying that “adakitic” features might be used as a local‐regional explo‐
ration tool for porphyry‐related ore deposits. The term "adakite" sensu stricto refers to a special geochemical composition of island arc magmas indicative of parental melt evolution outside the stability field of plagioclase, and melt equilibra‐
tion with residual garnet (Defant & Drummond 1990). Defant & Drummond (1990) argue that the appropriate P‐T conditions to stabilize or destabi‐
lize these mineral phases in an island arc setting associated with a thin layer of continental crust would apply to slab melting. Modern island arc subduction zones are usually characterized by low geothermal gradients such that the downgo‐
ing slab dehydrates before it melts; augmenting the local geothermal gradient sufficiently to facili‐
tate slab melting prior to significant slab dehydra‐
tion may only apply to a number of special geo‐
dynamic settings such as subduction zone initia‐
tion, subduction of young, hot oceanic litho‐
sphere (<25 Ma), or, possibly, flat subduction (e.g., Defant & Drummond 1990; Gutscher et al. 2000). The empirical observation of Thieblémont et al. (1997) might imply that special geodynamic set‐
tings facilitating slab melting (of which “adakitic” chemical compositions may be indicative of, but see below) favor the formation of porphyry‐
related ore deposits. As slab melting may be a highly effective means of oxidizing the upper mantle in the supraslab region, a positive correla‐
tion between slab melting and porphyry‐related ore deposit formation had also been envisaged from theoretical considerations based on ther‐
modynamic modeling (Mungall 2002). Indeed, Oyarzun et al. (2001) speculate that large por‐
phyry‐related ore deposits (partly associated with "adakitic" intrusions) of Late Eocene to Early Oli‐
gocene age in Chile formed in a special geody‐
namic setting (flat subduction) allowing slab melting to take place; in contrast, non‐“adakitic” Paleocene to Early Eocene porphyry‐related ore deposits in Chile are smaller in tonnage, and might be related to "standard" arc magmatism where slab melting does not contribute to arc magma genesis (Oyarzun et al. 2001). 2 Since the works of Thieblémont et al. (1997) and Oyarzun et al. (2001), the notion to associate "adakitic" magmatism with porphyry‐related ore deposits has gained some appeal in the economic geology literature (see review by Richards & Ker‐
rich 2007, and references therein). However, the conclusions of Oyarzun et al. (2001) have been vigorously debated (Oyarzun et al. 2002; Rabbia et al. 2002; Richards 2002; see also Richards & Kerrich 2007), in part because a straightforward "adakite"‐slab melt correlation cannot be unam‐
biguously demonstrated in a continental arc set‐
ting where similar chemical signatures (referred to as adakite‐like) may also be acquired through crustal magma evolution (Richards & Kerrich 2007). Moreover, even in island arc settings ada‐
kite‐like features of arc magmas may be pro‐
duced by other processes than slab melting, ei‐
ther in the mantle wedge (e.g., Castillo et al. 1999) and/or in the crust (Alonso‐Perez et al. 2009). Consequently, the (occasionally) observed association of adakite‐like magma chemistry and porphyry‐related ore deposits does not necessar‐
ily relate to a specific process in the magma source, but may also reflect certain petrogenetic processes of crustal magma evolution such as progressive volatile enrichment, which are equally regarded as favorable for porphyry‐style mineralization (Sillitoe 2000; Rohrlach & Loucks 2005; Richards & Kerrich 2007). The Late Tertiary Ecuadorian arc system at the NW South American margin hosts a number of moderate‐tonnage porphyry‐related ore deposits (Prodeminca 2000a, b), and regionally connects with the central‐northern Peruvian Tertiary arc whose Miocene metallogenic belt is of major economic importance (e.g., Noble & McKee 1999). Porphyry‐related ore deposits along the Ecuadorian margin formed on the background of a geodynamic setting characterized by a high de‐
gree of complexity involving Late Cretaceous oceanic terrane collision and major strike‐slip partitioning in the upper plate (e.g., Vallejo et al. 2006), the Late Oligocene break‐up of the Faral‐
lon plate and subsequent subduction of newly formed Nazca seafloor, abandoned spreading centers, and oceanic fracture zones at the Ecua‐
dorian trench (e.g., Lonsdale 2005), and collision of the Carnegie Ridge seamount chain with the margin since the Late Miocene (e.g., Michaud et al. 2009). The overall character of arc magmatism seems to shift from dominantly non‐adakitic to adakite‐like in the Late Miocene, possibly in re‐
sponse to the changing geodynamic regime, al‐
though the timing of the change is only loosely constrained (Chiaradia et al. 2004, 2009). The highly dynamic nature of the Tertiary Ecua‐
dorian margin represents an opportunity to study the complex interactions between geodynamic setting, arc magmatism, and ore deposit forma‐
tion. The principal aim of this thesis is to evaluate the mutual relevance of these different factors for each other. As the amount of available state‐
of‐the‐art geochronologic and geochemical data (outside of this thesis) on Tertiary arc magmatism and ore deposits in Ecuador is very limited, the present work can only serve as a first step to‐
wards a better understanding of the Tertiary Ec‐
uadorian arc magmatic evolution. In detail, this thesis aims to address the following issues: (1) A robust, regional geochronologic framework (as opposed to potentially disturbed ages re‐
lying on the K‐Ar isotopic system) for the Ec‐
uadorian arc segment and its porphyry sys‐
tems has not been established yet. This work presents robust zircon and titanite U‐Pb (ID‐
TIMS, LA‐MC‐ICP‐MS), molybdenite Re‐Os, and biotite as well as alunite 40Ar/39Ar (results were pending during thesis compilation, but will be available for subsequent manuscript editing) mass spectrometric data allowing the dating of intrusive and hydrothermal pulses of several Ecuadorian porphyry systems on a regional scale. These data are presented in Chapters 2 (magmatism) and 3 (hydrothermal systems). (2) Building on a pilot study by Chiaradia et al. (20004), this thesis provides new and extends existing datasets on Tertiary porphyry‐related arc magma isotopic compositions (Sr, Nd, Pb; Chapter 4) and geochemistry (multi‐element XRF, LA‐ICP‐MS; Chapter 5). These data may be used to track heterogeneous crustal basement units at depth, and to broadly con‐
strain the petrogenetic features of Tertiary arc magmatism. Particular emphasis is placed on the generation of adakite‐like features, and on their significance for porphyry‐related ore deposits. (3) Where appropriate in a given chapter of this thesis, I test spatio‐temporal correlations of ore deposit formation, variations in arc magma chemistry, and geodynamic changes at the Late Tertiary Ecuadorian margin to identify potential feedback mechanisms. This includes a critical review of the inferred geo‐
dynamic evolution of the Tertiary Ecuadorian arc system, and an adjustment of published convergence parameters to the Ecuadorian margin (parts of chapters 2 and 3; see also Appendix of Chapter 2). A short synthesis of the results of this analysis is presented in Chapter 6. References Alonso‐Perez, R., Müntener, O., Ulmer, P. (2009): Ig‐
neous garnet and amphibole fractionation in the roots of island arcs: experimental constraints on andesitic liquids. Contributions to Mineralogy and Petrology 157; 541‐558. Chiaradia, M., Fontboté, L., Beate, B. (2004): Cenozoic continental arc magmatism and associated mineraliza‐
tion in Ecuador. Mineralium Deposita 39; 204–222. Chiaradia, M., Müntener, O., Beate, B., Fontignie, D. (2009): Adakite‐like volcanism of Ecuador: lower crust magmatic evolution and recycling. Contributions to Mineralogy and Petrology 158; 563‐588. Cline, J. S. & Bodnar, R. J. (1991): Can economic por‐
phyry copper mineralization be generated by a typical calc‐alkaline melt? Journal of Geophysical Research 96, 8113–8126. Cooke, D. R., Hollings, P., Walshe, J. L. (2005): Giant porphyry deposits: characteristics, distribution, and tectonic controls. Economic Geology 100; 801‐818. Defant, M. J. & Drummond, M. S. (1990): Derivation of some modern arc magmas by melting of young sub‐
ducted lithosphere. Nature 347; 662‐665. Gutscher, M.‐A., Maury, R., Eissen, J.‐P., Bourdon, E. (2000): Can slab melting be caused by flat subduction? Geology 28; 535‐538. Jugo, P. J., Luth, R. W., Richards, J. P. (2005): An ex‐
perimental study of the sulfur content in basaltic melts saturated with immiscible sulfide or sulfate liquids at 1300°C and 1.0 GPa. Journal of Petrology 46; 783‐798. 3
Lonsdale, P. (2005): Creation of the Cocos and Nazca plates by fission of the Farallon Plate: Tectonophysics, v. 404, p. 237‐264. Richards, J. P. (2003): Tectono‐Magmatic Precursors for Porphyry Cu‐(Mo‐Au) Deposit Formation. Ec Geol. 98; 1515‐1533. Michaud, F., Witt, C., Royer, J. Y. (2009): Influence of the subduction of the Carnegie volcanic ridge on Ec‐
uadorian geology: reality and fiction. In: Kay, S. M., Ramos, V. A., Dickinson, W. R. (eds.), Backbone of the Americas: shallow subduction, plateau uplift, and ridge and terrane collision. Geological Society of America Memoir 204; doi: 10.1130/2009.1204(10). Richards, J. P. & Kerrich, R. (2007): Adakite‐like rocks: their diverse origins and questionable role in metal‐
logenesis. Economic Geology 102; 537‐376. Mungall, J. E. (2002): Roasting the mantle: slab melting and the genesis of major Au and Au‐rich Cu deposits. Geology 30; 915‐918. Noble, D. C. & McKee, E. H. (1999): The Miocene met‐
allogenic belt of central and northern Perú. SEG Spe‐
cial Publication 7; 155‐193. Oyarzun, R., Marquez, A., Lillo, J., Lopez, I., Rivera, S. (2001): Giant versus small porphyry copper deposits of Cenozoic age in northern Chile: adakitic versus normal calc‐alkaline magmatism. Mineralium Deposita 36; 794‐798. Oyarzun, R., Marquez, A., Lillo, J., Lopez, I., Rivera, S. (2002): Reply to Discussion on "Giant versus small porphyry copper deposits of Cenozoic age in northern Chile: adakitic versus normal calc‐alkaline magmatism" by Oyarzun et al. (Mineralium Deposita 36; 794‐798, 2001). Mineralium Deposita 37; 791‐794. Prodeminca (2000a) Evaluacion de distritos mineros del Ecuador, vol 2—Depositos epitermales en la Cor‐
dillera Andina. UCP Prodeminca Proyecto MEM BIRF 36–55 EC, Quito, Ecuador Prodeminca (2000b) Evaluacion de distritos mineros del Ecuador, vol 4—Depositos porfidicos y epi‐
mesotermales relacionados con intrusiones de las Cordilleras Occiental y Real. UCP Prodeminca Proyecto MEM BIRF 36–55 EC, Quito, Ecuador Rabbia, O. M., Hernandez, L. B., King, R. W., Lopez‐
Escobar, L. (2002): Discussion on "Giant versus small porphyry copper deposits of Cenozoic age in northern Chile: adakitic versus normal calc‐alkaline magmatism" by Oyarzun et al. (Mineralium Deposita 36; 794‐798, 2001). Mineralium Deposita 37; 791‐794. Richards, J. P. (2002): Discussion on "Giant versus small porphyry copper deposits of Cenozoic age in northern Chile: adakitic versus normal calc‐alkaline magmatism" by Oyarzun et al. (Mineralium Deposita 36; 794‐798, 2001). Mineralium Deposita 37; 788‐790. 4 Rohrlach, B. D. & Loucks, R. R. (2005): Multi‐million‐
year cyclic ramp‐up of volatiles in a lower crustal magma reservoir trapped below the Tampakan cop‐
per‐gold deposit by Mio‐Pliocene crustal compression in the southern Philippines. In: Porter, T. M. (ed.), Su‐
per Porphyry Copper & Gold Deposits: A Global Per‐
spective; PGC Publishing, Adelaide, v. 2; 369‐407. Rosenbaum, G., Giles, D., Saxon, M., Betts, P.G., Weinberg, R.F., Duboz, C. (2005): Subduction of the Nazca Ridge and the Inca Plateau: insights into the formation of ore deposits in Peru. Earth and Planetary Science Letters 239; 18–32. Seedorff, E., Dilles, J. H., Proffett, J. M. Jr., Einaudi, M. T., Zurcher, L., Stavast, W. J. A., Johnson, D. A., Barton, M. D. (2005): Porphyry deposits: characteristics and origin of hypogene features. Economic Geology 100th Anniversary Volume; 251‐298. Sillitoe, R. H. (2000): Gold‐rich porphyry deposits: de‐
scriptive and genetic models and their role in explora‐
tion and discovery. Society of Economic Geologists Reviews 13; 315 – 345. Thiéblemont, D., Stein, G., Lescuyer, J.‐L. (1997): Gisements épithermaux et porphyriques: la connexion adakite. C. R. Academy of Sciences, Paris, Sciences de la terre et des planets/Earth and Planetary Sciences 325; 103‐109. Tosdal, R.M. & Richards, J. P. (2001): Magmatic and structural controls on the development of porphyry Cu±Mo±Au deposits. Reviews in Economic Geology 14; 157–181. Vallejo, C., Spikings, R.A., Luzieux, L., Winkler, W., Chew, D., Page, L., (2006): The early interaction be‐
tween the Caribbean Plateau and the NW South American Plate. Terra Nova 18, 264–269 Wilkinson, B. H. & Kesler, S. E. (2009): Quantitative identification of metallogenic epochs and provinces: application to Phanerozoic porphyry copper deposits. Economic Geology 104; 607‐622. CHAPTER II GEODYNAMIC CONTROLS ON TERTIARY ARC MAGMATISM IN ECUA‐
DOR: CONSTRAINTS FROM U‐Pb ZIRCON GEOCHRONOLOGY OF OLI‐
GOCENE‐MIOCENE INTRUSIONS AND REGIONAL AGE DISTRIBUTION TRENDS Abstract We obtained U‐Pb zircon ages of Late Tertiary intrusions in the northern Andes to provide robust time cali‐
bration points for the intrusive geochronologic framework of Ecuador which is mostly based on K‐Ar data. Intrusion emplacement ages range from about 31 to 7 Ma, and mainly pool in the Late Oligocene‐Early Miocene. Where both K‐Ar and U‐Pb data exist for a given intrusive system, ages obtained by the different methods are usually concordant within 1‐4 m.y. implying that K‐Ar ages may be used as proxies for the tim‐
ing of Tertiary arc magmatism on a regional scale. Except for a single sample, the investigated intrusions completely lack externally inherited zircons, in agreement with dominantly zircon‐poor, oceanic crustal basement domains. Spatio‐temporal distribution trends of Tertiary arc magmatism inferred from screened U‐Pb, K‐Ar, and zir‐
con fission track geochronologic data allow tracking of the progressive broadening of a flat slab region be‐
low southernmost Ecuador in the Mid‐ to Late Miocene, and of moderate slab shallowing in northern‐
central Ecuador in the Late Miocene. These regional arc migration patterns correlate in time with the sub‐
duction of the buoyant Inca Plateau and the Carnegie Ridge seamount chain. The temporal distribution of Tertiary Ecuadorian arc magmatism indicates a Late Oligocene‐Early Miocene arc magmatic flare‐up event comprising widespread ignimbrite eruption and batholith construction. Initiation of the flare‐up event coin‐
cides in time with accelerating, less oblique Farallon/Nazca‐South America plate convergence, suggesting a positive feedback between convergence rates, asthenospheric melt production, mantle‐crust melt flux, and upper crustal arc magmatic productivity in Ecuador. 5
Introduction The spatio‐temporal distribution of arc magma‐
tism, typically manifested as distinct belt seg‐
ments at the Earth's surface, is controlled by the interplay of multiple tectonomagmatic parame‐
ters which derive from the complex interactions of descending slab, overriding plate, and the mantle wedge (e.g., Hamilton 1995). A number of studies evaluate feedback processes between the geodynamic evolution and subduction‐related Mesozoic‐Cenozoic arc magmatism along the western plate edge of South America (e.g. Jaillard & Soler 1996; Kay et al. 2005), but data for the Late Tertiary tectonomagmatic evolution of the northern Andes are lacking. In this contribution, we are presenting the first dataset of robust U‐Pb zircon ages of Late Tertiary intrusions in the northern Andes of Ecuador which to date have only been characterized by K‐Ar geochronologic data. Our new data, combined with previously pub‐
lished radiometric age information, allow us to discuss the relationships between arc magmatism and changes in the Tertiary geodynamic regime of the northern Andean margin. Because a side outcome of this study (discussed in Chapters 3 and 5) is to investigate links between the geody‐
namic setting, magma chemistry, and intrusion‐
related mineralization in Ecuador, our study fo‐
cuses on intrusions spatially associated with por‐
phyry‐related ore deposits and their respective host rocks. Mineralized porphyry intrusions are typically small (few km2 outcrop area), but are thought to represent the uppermost crustal manifestations of significantly larger magmatic systems at depth, which constitute the major source of the mineralizing fluids (e.g., Sillitoe 1973). Therefore, knowledge of the age of em‐
placement of porphyry stocks can serve as a proxy for the timing of more voluminous, not yet unroofed plutonism at depth. Regional Geology of Ecuador Ecuador’s fundamental physiographic elements coincide with major geologic domains and com‐
prise the flat‐lying western Costa forearc and eastern Oriente foreland regions, enclosing be‐
tween them the central Andean chain which splits into the Western and Eastern Cordillera, separated by the Interandean Depression (IAD; Litherland et al. 1994; Fig. 1). Allochthonous ma‐
fic‐ultramafic oceanic basement domains, inter‐
preted as hotspot‐derived oceanic plateau frag‐
ments and mainly accreted in the Late Creta‐
ceous, floor the present‐day forearc and frontal arc regions of Ecuador (e.g., Vallejo et al. 2009). Locally, especially in the Western Cordillera and its western foothills, these oceanic plateau units host or are tectonically juxtaposed against sev‐
eral pre‐ and post‐accretionary island arc com‐
plexes of Late Cretaceous‐Early Tertiary age (Vallejo et al. 2009). The present‐day main arc is built upon IAD basement units which are likely heterogeneous in nature, comprising tectonized slices of Eastern Cordillera and oceanic plateau material (Feininger & Seguin 1983; Spikings et al. 2005; Chiaradia et al. 2009). Older Tertiary subaerial arc volcanic formations are exposed in the Interandean region south of 2.5°S where ac‐
tive volcanism of the Northern Volcanic Zone ceases (Hungerbühler et al. 2002). Characteristics of the major Tertiary arc volcanic formations of Ecuador are listed in Table 1. Two prominent N‐ to NNE‐trending regional fault zones structure the Western Cordillera (Fig. 2). The Calacalí‐Pujili‐Pallatanga fault zone (CPPF) forms the Cordillera's eastern structural limit; it intersects the Western Cordillera at 3°S to splay off towards the Gulf of Guayaquil (Winkler et al. 2005). The Chimbo‐Toachi shear zone (CTSZ) forms the eastern limit of the Macuchi island arc Figure 1 (next page): Simplified geological map of the Cordillera region of Ecuador, focusing on Tertiary arc magmatic units; magmatic centers investigated in this study are marked. Inset shows present‐day geodynamic situation of the Ecuadorian‐Colombian margin and the Panama basin; dark gray areas outline seamount chains (offshore) and Cordil‐
lera ranges (onshore); seafloor features include active spreading centers (thick lines), extinct spreading centers (thin
lines), active faults (hairlines), and transform faults/scarps (dashed hairlines). Inset adapted from Meschede & Barck‐
hausen (2001); main map adapted from Litherland et al. (1994), Steinmann (1997), Dunkley & Gaibor (1997), McCourt
et al. (1997), Pratt et al. (1997), Hughes et al. (1998), and Palacios et al. (2008). 6 7
sequence (Hughes & Pilatasig 2002). Seismic studies suggest that, while subvertical in their uppermost portions, both fault zones extend down to mid‐deep crustal levels where they are defined by 35° E‐dipping fault planes (Guillier et al. 2001). Tertiary intrusions of the Western Cor‐
dillera are aligned along the CTSZ and, towards its southern termination, the CPPF (Fig. 1, 2). In SW Ecuador the Amotape terrane forms a distinct tectonic unit which extends further south into northern Peru (Mitouard et al. 1990; Litherland et al. 1994). Its Paleozoic basement units, ex‐
posed in the El Oro massif in Ecuador, are mostly covered by volcanic‐sedimentary sequences of the Cretaceous Celica‐Lancones basin (Jaillard et al. 1996). The El Oro massif, interpreted as a mi‐
crocontinental block accreted to South America in the earliest Cretaceous (Litherland et al. 1994), represents a major structural break from the main Andean strike further north in that it adds a prominent ESE‐WNW structural trend to the re‐
gional tectonic framework which structurally fo‐
cused intrusion emplacement in across‐arc di‐
mension (Figs. 1, 2). Assembly of multiple exotic terrane fragments, syn‐ and post‐accretionary block rotation and fragmentation, large‐scale forearc sliver dis‐
placement, and an oblique subduction setting, concentrated in a narrow range of latitudes, pro‐
duced a tectonically heterogeneous Tertiary arc system in Ecuador (e.g., Mitouard et al. 1990; Hungerbühler et al. 2002; Spikings et al. 2005; Vallejo et al. 2009). Therefore, it is useful to in‐
spect the arc in its regional context along the NW South American margin. Based on their spatial distribution and tectonic context, three major Tertiary arc segments (northern, central, and southern Ecuador) can be distinguished in Ecua‐
dor (Fig. 2). The regional outcrop pattern of Ter‐
tiary arc magmatic units of NW South America demonstrates an essentially uninterrupted mag‐
matic chain developed along the northern Peru‐
vian (Calipuy) and southern Ecuadorian margin, geometrically continuous with the central‐
northern Ecuadorian arc segments. The north‐
wards continuation of the northern arc segment into southern Colombia is less obvious, as this area, the Naranjal block, is characterized by a high degree of structural complexity including the appearance of additional Late Cretaceous island 8 arc units (Rio Cala arc; Vallejo 2007; Chiaradia 2009). The Tertiary arc cuts across a fundamental base‐
ment contrast in the Saraguro arc segment, the exact location of which is not known (cf. Chapter 4). It might be located close to the CPPF at ca. 3°S, or further south towards the Jubones or Pi‐
ñas‐Portovelo faults, and corresponds to the northern limit of Amotape basement units vs. allochthonous Cretaceous oceanic plateau mate‐
rial (Litherland et al. 1994; Pratt et al. 1997; Spik‐
ings et al. 2005). Petrologically, most pre‐
Miocene arc units north of 2.5‐3°S and west of the CPPF represent a progressively maturing is‐
land arc system emplaced in a submarine envi‐
ronment on oceanic plateau basement. In con‐
trast, subaerial arc magmatism taking place east of the CPPF and in southern‐central Ecuador pro‐
duced typical continental arc sequences. Tertiary arc segmentation and geology of investigated magmatic centers The northern Ecuadorian arc segment The northern Ecuadorian arc segment comprises the whole Tertiary arc system of central‐northern Ecuador (Fig. 2). Its main outcrop unit is the Pa‐
leocene‐Eocene Macuchi Unit, representing a submarine sequence of pillow lavas and hyalo‐
clastites, and their redistributed sedimentary equivalents (Hughes & Pilatasig 2002; Vallejo 2007). Due to deep erosion levels in the Western Cordillera, post‐Macuchi arc volcanics are spa‐
tially underrepresented. Aerially extensive Oligo‐
cene‐Miocene volcanic cover sequences can be inferred from prominent coeval batholith intru‐
sions in central and northern Ecuador (see be‐
low). Locally preserved post‐Macuchi volcanic or volcaniclastic rocks comprise: (1) minor Saraguro Group volcanics cropping out in small, isolated patches in the central Western Cordillera close to the CPPF. In southern Ecuador, the Saraguro Group is mainly of Oligocene‐Early Miocene age (Hungerbühler et al. 2002; Tab. 1), but Saraguro Group (Ocaña Formation) volcanics overlying the Paleogene Yunguilla turbidites at the eastern edge of the Western Cordillera have K‐Ar and 9
zircon fission track (ZFT) ages of 36‐39 Ma, thus overlapping in age with the youngest Macuchi Unit (Dunkley & Gaibor 1997; Tab. 1); (2) the Oli‐
gocene‐Early Miocene San Juan de Lachas Forma‐
tion, contemporaneous with the Saraguro Group, crops out in small areas close to the Ecuadorian‐
Colombian border, where it overlies the Macuchi Unit (Vallejo 2007); (3) Miocene volcaniclastic Figure 2: Regional Tertiary arc outcrop pattern along the NW South American margin showing arc seg‐
ments as discussed in the text (Northern, Central, and Southern Ecuador, and Calipuy, Peru), and major fault systems which con‐
trolled Tertiary intrusion emplacement in Ecuador. Tertiary intrusions (italics) are aligned along major structures (thick lines) or inferred major structures (dashed thick lines). The Raspas complex of the NW El Oro range indicates prox‐
imity to the ancient Amo‐
tape suture zone (Bosch et al. 2002). Age ranges of vol‐
canic‐volcaniclastic forma‐
tions referenced in the leg‐
end are approximate, and refer to the base of the cor‐
responding unit; Calipuy Group volcanism extends throughout most of the Ter‐
tiary (Navarro et al. 2008). Also shown are Wadati‐
Benioff zones (Gutscher et al. 1999a, b; Guillier et al. 2001) outlining the flat slab segment below northern Peru and southern Ecuador. Map compiled from same sources as Fig. 1, plus Gómez Tapias et al. (2006), and Winkler et al. (2005). 10 rocks of the Zumbagua Unit in central Ecuador unconformably overlie Eocene turbidites (Hughes et al. 1998); (4) Pliocene‐Holocene volcanism in Ecuador and Colombia, constituting the present‐
day Northern Volcanic Zone of the Andes, covers part of the older arc units of the Western Cordil‐
lera and is prominently exposed further east‐
wards in the IAD. Considering the relatively narrow width of the Macuchi arc segment to the west of the CPPF (mostly 50 km or less) com‐
pared to a typical average arc width of 97 km (Stern 2002), one expects Pliocene‐recent volcan‐
ics to conceal the landward extent of Miocene or older volcanic sequences spatially associated with the northern Ecuadorian arc segment which, for example, might constitute the source area for the Paleocene Silante or the Miocene Zumbagua vol‐
caniclastic units of the Western Cordillera. Large, mainly Macuchi Unit‐hosted intrusions are unroofed in Ecuador's Western Cordillera and its foothills. Relatively few, mostly small‐sized Eo‐
cene stocks and plutons are exposed, with the Santiago pluton close to the Ecuadorian‐
Colombian border representing the only major Eocene plutonic body (Boland et al. 1998). In con‐
trast, several large Oligocene‐Miocene batholiths form two regional intrusive clusters and were investigated in this study. In northern Ecuador the Junin and Cuellaje por‐
phyry systems are emplaced in the Apuela‐
Nanegal batholith, which in turn is hosted by vol‐
caniclastic units ascribed to the Late Cretaceous‐
Early Tertiary Rio Cala and Macuchi island arc se‐
quences (Boland et al. 1998; Prodeminca 2000a; Chiaradia 2009). The Apuela‐Nanegal batholith covers an area of 750 km2 thus forming one of the largest Tertiary intrusive complexes in Ecua‐
dor. It dominantly comprises hornblende‐ and biotite‐bearing quartz‐diorites and granodiorites which are punctured by several porphyry stocks and dikes of variable, mainly dacitic, composition. Available geochronological data (hornblende, biotite and whole rock K‐Ar; Van Thournout 1991; MMAJ/JICA 1992; Boland et al. 1998) indicate a period of protracted magmatism from 19 to 6 Ma, although some K‐Ar ages might reflect ther‐
mal resetting caused by intense hydrothermal alteration associated with porphyry stock em‐
placement. In northern‐central Ecuador the hornblende‐ and biotite‐bearing tonalitic‐granodioritic plutonic complexes of Telimbela‐Chazo Juan and Balsa‐
pamba‐Las Guardias host several porphyry intru‐
sions and associated hydrothermal systems; they are situated at a close distance to the western Echeandia‐La Industria and the northern Corazon plutons (McCourt et al. 1997; Prodeminca 2000a). Combined, these intrusive complexes are of batholithic dimension and form the most prominent cluster of plutons exposed in western central Ecuador. Available K‐Ar ages (hornblende or biotite) for the Telimbela‐Chazo Juan pluton range from 21‐15 Ma whereas ages for the Balsa‐
pamba‐Las Guardias pluton are mostly older (34‐
30 Ma) except for two hornblende K‐Ar ages of 19.8±3.0 Ma and 25.7±0.9 Ma (Kennerley 1980; MMAJ/JICA 1989, 1991; McCourt et al. 1997). Furthermore, the hydrothermal system of a por‐
phyry intrusion hosted by the Balsapamba pluton has been dated at 19.7 ± 0.3 Ma (Re‐Os molyb‐
denite; Chiaradia et al. 2004). Ages reported for the nearby plutons of Echeandia‐La Industria (27‐
23 Ma) and Corazon (16‐14 Ma) closely correlate in time with these ages (Kennerley 1980; MMAJ/JICA 1989, 1991; McCourt et al. 1997). The central Ecuadorian arc segment The central Ecuadorian arc segment, structurally bracketed between the Piñas‐Portovelo Fault at c. 3° 45' S, and the CPPF at c. 2° 10' S (Fig. 2) hosts mainly subaerial volcanic formations spanning the whole Tertiary (Tab. 1), with the Oligocene‐
Early Miocene Saraguro Group forming the major outcrop unit. Several Miocene basins filled with volcaniclastic‐sedimentary sequences overlie the Saraguro Group volcanics and form prominent outcrop units in this arc segment (Steinmann 1997; Hungerbühler et al. 2002). The oldest volcanic formation is the Paleocene (?) Sacapalca Formation which mainly occurs in the northwards projection of the Catamayo graben (Hungerbühler et al. 2002), but also well to the west of it (Pratt et al. 1997). The Eocene Chinchín Formation forms the base of the local Quingeo basin SW of Cuenca (Steinmann 1997; Hunger‐
bühler et al. 2002). Post‐Saraguro volcanism is represented by the Sta. Isabel and Quimsacocha formations which overlie older volcanics in the southwestern prolongation of the Cuenca basin. The Late Miocene‐Pliocene Tarqui Formation represents the youngest major volcanic forma‐
tion in southern Ecuador, with outcrop areas concentrated along the eastern border of the Saraguro arc segment, where it overlies Saca‐
11
palca, Chinchín, and Saraguro volcanics (Pratt et al. 1997; Hungerbühler et al. 2002). Two major and one minor intrusive centers of this arc segment were investigated in this study. Tonalite, quartz‐diorite, and granodiorite units of the Chaucha batholith, exposed adjacent to the CPPF, intrude or are tectonically juxtaposed against undifferentiated metasedimentary rocks, mafic lavas of the Pallatanga Unit, sedimentary rocks of the Late Cretaceous‐Eocene Yunguilla and Angamarca turbidite series, and Saraguro Group volcanics (Dunkley & Gaibor 1997; Prodeminca 2000a). The Chaucha batholith hosts several porphyry intrusions and its SE portion has been dated at 13.3‐9.8 Ma (K‐Ar various minerals; Kennerley 1980; INEMIN‐AGCD 1989; additional references in Prodeminca 2000a). Along the northern limit of the Amotape range a series of mainly diorite, quartz‐diorite and grano‐
diorite intrusions forms a WNW‐ESE trending belt hereafter referred to as Cangrejos‐Zaruma intru‐
sive belt. These plutons intrude the southern flank and partly the hinge of a WNW‐ESE trending regional antiform. Amongst others, the plutonic complexes comprise, from west to east, Cangre‐
jos, Paccha, and El Poglio, and are associated with several mineralized porphyry intrusions as well as a large epithermal vein system at Zaruma and Portovelo (Pratt et al. 1997; Spencer et al. 2002). The regional basement is mostly obscured by overlying Saraguro Group volcanics which form the major host lithology for the Tertiary intru‐
sions, but it is exposed as several metamorphic inliers along the intrusive belt. Towards the east the belt is disrupted by a network of Andean‐
trending (NNE) regional faults, which juxtapose units of the Saraguro Group against the Sacapalca Formation further east (Pratt et al. 1997). Spencer et al. (2002) regard the Sacapalca‐hosted Fierro Urcu porphyry center as the eastern pro‐
longation of the Cangrejos‐Zaruma intrusive belt. The belt is inferred to be mainly Miocene in age but only a single K‐Ar age of 16.9 ± 0.2 Ma for the Paccha intrusion has been reported so far (Pratt et al. 1997). The Paleocene‐Early Eocene San Lu‐
cas pluton (66‐52 Ma; Aspden et al. 1992), occur‐
ring at the eastern termination of the Cangrejos‐
Zaruma intrusive belt at the border to the Saca‐
palca arc segment (Fig. 2), significantly predates western belt plutonism. 12 In a tectonized zone characterized by regional CPPF splay faults, the Gaby‐Papa Grande por‐
phyry system forms a minor intrusive complex. The porphyry system is hosted by basalts of the Pallatanga oceanic plateau unit and consists of several small (few km2 in total) intrusive units comprising a tonalite pluton as well as plagio‐
clase‐hornblende porphyry stocks and dikes. Prodeminca (2000a) report a K‐Ar age of 19.3 ± 1.0 Ma for the Gaby porphyry without providing further details as to which mineral was used for dating. Apart from these deeply eroded intrusive cen‐
ters, a number of subvolcanic domes and plugs associated with epithermal ore deposits were investigated in this work. These comprise the high sulfidation type epithermal systems of Quimsacocha and El Mozo, as well as the poly‐
metallic vein deposit of Tres Chorreras; spatially associated subvolcanics are described or inferred for each of these deposits (Prodeminca 2000b). Quimsacocha forms a topographically prominent caldera of c. 5 km diameter, surrounded by coge‐
netic andesitic lavas and breccias, and rhyolitic ignimbrites (Beate et al. 2001). The caldera hosts several rhyodacitic domes whose emplacement postdates the epithermal mineralization event at the caldera flank. Zircon fission track ages ob‐
tained on ignimbrites (5.2‐4.9 Ma) and subvol‐
canic intra‐caldera domes (3.6 Ma) indicate a Late Miocene phase of volcanic activity (Beate et al. 2001). Polymetallic veins and mineralized breccia bodies associated with the Tres Chorreras diatreme complex, situated between the Chaucha, Gaby, and Quimsacocha magmatic cen‐
ters, are hosted by silicic volcanic units ascribed to the Saraguro Group (Pratt et al. 1997). Prodeminca (2000a) mentions the occurrence of syn‐ to postmineral intrusions, but we failed to clearly identify these in the field due to intense hydrothermal alteration. The El Mozo high sulfi‐
dation type epithermal mineralization occurs at the western flank of the Eastern Cordillera. The El Mozo complex includes strongly altered volcanic units which are hosted by rhyolitic tuffs of the uppermost Saraguro Group (La Paz Formation). Exploration drill cores encountered altered granodiorite porphyry intrusions occurring at shallow depth below the epithermal high 13
sulfidation mineralization; an alunite K‐Ar age of 15.4±0.7 Ma is thought to date the high‐
sulfidation alteration and mineralization event (Prodeminca 2000b). yielded a Late Cretaceous age. The Curiplaya por‐
phyry system is thus associated with magmatism of the Tangula batholith, rather than with Terti‐
ary Sacapalca‐Loma Blanca magmatism. The southern Ecuadorian arc segment Analytical techniques and sample material Outcrop units of this arc segment south of 3° 45' S are mainly preserved in the N‐S trending Cata‐
mayo graben at the border between the Eastern Cordillera and the Amotape terrane in southern Ecuador (Figs. 1, 2). An inferred Late Maas‐
trichtian‐Paleocene age of the volcanic Sacapalca Formation is poorly resolved and based on a sin‐
gle ZFT age of 67±6 Ma (Hungerbühler et al. 2002), in conjunction with the Paleocene‐
Miocene ages of several plutons intruding the Sacapalca Formation, including the major Paleo‐
cene‐Early Eocene San Lucas pluton at the border to the Saraguro arc segment (Aspden et al. 1992; Jaillard et al. 1996). Locally, volcanics of the Late Eocene‐Early Oligocene Loma Blanca Formation, forming the base of the Saraguro Group, overlie the Sacapalca Formation (Hungerbühler et al. 2002). The major Early to Mid‐Miocene Portachuela batholith is exposed in the Eastern Cordillera directly to the east of the Las Aradas fault (Aspden et al. 1992). In combination with several minor, Sacapalca‐hosted intrusions, it demonstrates Early to Mid‐Miocene arc magma‐
tism in this region where the young volcanic cover has not been preserved. We did not study any magmatic centers directly associated with this arc segment in the present work, but we investigated the Curiplaya porphyry system, which is located c. 60 km W of the Cata‐
mayo graben and c. 25 km W of the westernmost Loma Blanca and Sacapalca Formation outcrops. The Curiplaya porphyry intrusions are hosted by Cretaceous volcanic rocks of the Celica Formation in the Celica‐Lancones basin and crop out proxi‐
mal to the major Tangula batholith of Mid‐Late Cretaceous age (Hall & Calle 1982; Palacios et al. 2008). The quartz‐diorite porphyry intrusions were inferred to be Tertiary in age and are partly overlain by dacitic tuffs tentatively assigned to the Loma Blanca Formation. They are structurally disrupted by a complex network of NNE‐SSW and ESE‐WSW trending faults (Howe International 2006). However, anticipating our results, a Curiplaya porphyry intrusion dated in this study 14 We used both isotope dilution thermal ionization mass spectrometry (ID‐TIMS; in the following re‐
ferred to as TIMS) and laser ablation multi‐
collector inductively coupled plasma mass spec‐
trometry (LA‐MC‐ICP‐MS) for U‐Pb isotopic measurements to obtain age information on our samples. Samples where, based on the regional geological setting (i.e., proximity to basement units with a high potential for recycled crustal components; e.g., Noble et al. 1997; Vallejo 2007), we suspected significant external zircon inheritance were dated by LA‐MC‐ICP‐MS whereas all other samples were dated by TIMS. Samples were crushed and milled to <350 μm, and inclusion‐free zircons suited for isotopic analysis were handpicked from the nonmagnetic heavy mineral (>3.32 g/cm3) fraction using a bin‐
ocular. The obtained zircon fractions were either processed for annealing (TIMS) or mounted in epoxy and polished such that grain interiors were exposed for scanning electron microscopy and cathodoluminescence (SEM‐CL) imaging and, where applicable, LA‐MC‐ICP‐MS analysis. Sample material Sampling details are summarized in Table 2. Out‐
crop sampling for zircon dating comprised c. 5 kg of rock material; proximal lithologic contact zones or rock heterogeneities (xenoliths, en‐
claves) were avoided as far as possible. In addi‐
tion, a number of drill core samples were col‐
lected for this study; due to its limited availability drill core sample quantities are smaller (c. 1‐2 kg). Where not directly avoidable during sampling, rock heterogeneities were removed later using a diamond‐blade disc saw. For samples where TIMS single grain analyses were performed we indi‐
vidually imaged ca. 20‐30 zircons/sample by SEM‐
CL. For zircons where U‐Pb measurements were carried out by means of LA‐MC‐ICP‐MS only bulk sample (25‐50 zircons) SEM‐CL images were ob‐
tained which do not provide sufficient resolution to discuss zircon textures in detail. Zircon SEM‐CL characteristics are summarized in Table 2 and a selection of SEM‐CL images is presented in Figure 3. TIMS analysis Zircon single grain (and, in one case, a two‐grain fraction) dissolution and U‐Pb separation at the Department of Mineralogy, University of Geneva, followed the techniques described by Ovtcharova et al. (2006). An initial annealing/leaching "chemical abrasion" zircon treatment step (Ov‐
tcharova et al. 2006, based on Mattinson 2005) served to minimize effects of post‐
crystallizational radiogenic Pb loss which are oth‐
erwise expected to be significant given the high degree of zircon‐fluid interaction displayed by some zircons (Fig. 3). Samples were spiked using a mixed 205Pb‐233U‐235U spike solution, and zircon dissolution in 63 μl concentrated HF with a trace of 7N HNO3 took place at 180°C for seven days, followed by evaporation and overnight re‐
dissolution in 36 μl 3N HCl. Isotopic analyses at the University of Geneva were performed using a Thermo Fisher TRITON mass spectrometer equipped with a MasCom‐2 electron multiplier and a digital ion counting sys‐
tem. Loading of U and Pb on previously out‐
gassed single Re filaments took place using 1 μl of a silica gel‐H3PO4 mixture (Gerstenberger & Haase 1997). Lead isotopes were measured by peak‐hopping on the MasCom‐2 electron multi‐
plier, and U isotopes as oxides were measured either by peak‐hopping on the MasCom‐2 elec‐
tron multiplier or, at signal intensities of >3 mV, simultaneously (static mode) on Faraday cups linked to amplifiers equipped with 1012 Ohm re‐
sistors. Mass fractionation of Pb (0.08±0.05%/amu) was controlled by SRM‐981 standard measurements. Mass fractionation of U was corrected online by using a double 233U‐235U spike solution. The total procedural common lead blank was 2.07 ± 1.97 pg (average of 20 total blank meas‐
urements in the 2007‐2008 period) and has the following isotopic composition (at 2σ uncertainty, fractionation‐corrected): 206/204Pb: 18.36±0.34; 207/204
Pb: 15.59±0.20; 208/204Pb: 38.00±0.69; the total blank isotopic composition did not vary sys‐
tematically over the range of total blank common Pb amounts (0.5‐7.4 pg). Following the initial an‐
nealing/leaching step sample common Pb con‐
tents were identical within error to total blank common Pb amounts and were thus solely at‐
tributed to laboratory contamination except for a minor number of samples where part (<1.2 pg) of the common Pb was corrected with the isotopic composition of Stacey & Kramers (1975) using an appropriate sample age estimate based on geo‐
logical field relationships and published geochro‐
nologic data. Whole rock Th and U contents measured by multi‐element ICP‐MS analysis (Chapter 5) were used as a proxy to estimate melt Th/U ratios for Th disequilibrium correction. Estimated Th/U ratios are typically in the range of 2‐3, and arbitrary modification of this ratio to up to four (a commonly assumed Th/U ratio if geo‐
chemical information is lacking; e.g., Ovtcharova et al. 2006) does not change the obtained U‐Pb ages beyond analytical uncertainties. The uncertainties of spike and blank Pb isotopic composition, mass fractionation correction, and tracer calibration were propagated to the final uncertainties of isotopic ratios and ages of each individual analysis. In addition, uncertainties in the decay constants of 238U and 235U (238U: 0.16%, 235
U: 0.21%; Jaffey et al. 1971 values with an ad‐
ditional uncertainty factor of 1.5 as suggested by Mattinson (1987) to ensure direct compatibility with LA‐MC‐ICP‐MS data) were propagated sepa‐
rately and added quadratically to weighted mean or single zircon uncertainties discussed in the text. 207Pb/235U age information is only used to evaluate the concordancy of individual zircon analyses. Concordia plots and weighted average age calculations were prepared using the Isoplot v.3.31 Excel macro of Ludwig (2003). All uncer‐
tainties and error ellipses are reported as 2‐ , and weighted mean 206Pb/238U ages are pre‐
sented at 95% confidence level. LA‐MC‐ICP‐MS analysis Epoxy grain mounts of sample zircons and SL‐1 standard zircon fragments for controlling inter‐
element fractionation (TIMS age 563.5 ± 3.2 Ma; Gehrels et al. 2008) were prepared at the Univer‐
sity of Geneva, and isotopic measurements took place at the Arizona LaserChron Center, Univer‐
sity of Arizona. Zircons measured by 15
Figure 3: Zircon SEM‐CL images of samples documenting rounded core domains, incremental zircon growth stages, and variable degrees of zircon‐fluid interaction; white scale bar is 100 μm. Note that none of the zircons dated by
TIMS showed major external (i.e., xenocrystic) inherited age components. A – Continuous oscillatory zoning pattern (OZP) with resorbed low‐CL internal zone and concordant low‐CL rim; several melt inclusions disturb OZP pattern (E06140). B – Zircon interior with large re‐homogenized low‐CL domains and high‐CL fractures interpreted as altera‐
tion by zircon‐hydrothermal fluid interaction; thin overgrowth rim with well‐developed OZP. High‐CL fractures termi‐
nate against the fresh overgrowth rim possibly indicating alteration of the antecrystic core took place prior to zircon re‐immersion into the melt and final intrusion solifidication (E07032). C – Continuous OZP with minor resorption tex‐
tures (E05090). D – Multiple resorbed OZP domains and overgrowth zones (E06206). E – Zircon with partly resorbed and recrystallized internal domains, as well as conspicuous alteration of the outer rim (E07005). F – Well‐defined OZP domains which are slightly resorbed in places, and alteration along grain margins (E07018). 16 LA‐MC‐ICP‐MS were not annealed/leached prior to analysis and are thus potentially more likely to display Pb loss features than annealed/leached zircons measured by TIMS. However, in most cases the magnitude of this effect is probably smaller than the analytical LA‐MC‐ICP‐MS preci‐
sion, such that zircon ages obtained by LA‐MC‐
ICP‐MS are not supposed to be systematically biased outside of their analytical uncertainty range. Analytical procedures and measurement condi‐
tions for LA‐MC‐ICP‐MS analysis on a GV Instru‐
ments Isoprobe with an attached New Wave/Lambda Physik DUV193 Excimer laser are outlined in Gehrels et al. (2008). Laser ablation spot diameters were 25 or 35 μm. Measurements were carried out in static mode using Faraday detectors for 238U, 232Th, 208Pb, 207Pb and 206Pb, and an ion‐counting channel for 204Pb. Back‐
ground on‐peak measurement for 20s with the laser off was followed by 20 one‐second integra‐
tions with the laser on; delay time between two samples was 30s. The analyses were corrected for common Pb using the measured 204Pb and an as‐
sumed initial lead isotopic composition of Stacey & Kramers (1975). Random measurement errors at 2‐б level were propagated into individual analyses. As the used standard material is of significantly older age than the unknown zircons, relative random errors of unknowns are typically 2‐3 times higher than random errors of the standard zircon. Measured low 207Pb intensities (<0.4 mV) resulted in very large errors in the 207Pb/235U and 206Pb/207Pb ra‐
tios producing poorly reliable ages calculated from these isotopic ratios; we therefore exclu‐
sively use 206Pb/238U age information. Additional systematic errors, discussed in detail by Gehrels et al. (2008), comprise uncertainties of the U de‐
cay constants (cf. TIMS analytics), SL‐1 standard zircon age, fractionation correction, and common Pb correction. Systematic errors were propagated separately yielding an average 2‐б error of 1.42±0.54% on 206Pb/238U ages; this average error was added quadratically to final weighted mean age uncertainties. LA‐MC‐ICP‐MS data reduction The MSWD is usually used for population control of zircon LA‐MC‐ICP‐MS data: if the MSWD of all analyzed zircons significantly deviates from unity, histograms or cumulative probability plots are used to statistically evaluate the number of zir‐
con populations present. To explain polymodal age distributions investigators often invoke ra‐
diogenic Pb loss caused by zircon‐fluid interaction to reject young zircon ages (e.g., Maksaev et al. 2004), and zircon inheritance or ‘subtle inheri‐
tance’ to reject old zircon ages (e.g., Campbell et al. 2006). Zircon ages at the lower and upper age range are then progressively rejected until the MSWD is decreased to a statistically acceptable value, typically on the order of the analyzed standard zircons (e.g., Bryan et al. 2008). How‐
ever, unless a large number of zircons are ana‐
lyzed, the usage of histograms or cumulative probability plots, where inflection points are used to identify different zircon populations expected to follow a Gaussian distribution, can fail to clearly resolve multiple or mixed age populations. If the magnitude of assigned analytical random errors is correct, the expected MSWD value for a single zircon population following a Gaussian dis‐
tribution is always unity. Deviation from unity either indicates underestimation of analytical random errors or reflects real geological scatter which might be caused by analytically irresolv‐
able external inheritance, antecrystic compo‐
nents or Pb loss features. As outlined by Wendt & Carl (1991) a standard deviation can be assigned to the MSWD, the size of which depends on the system's degrees of freedom, f, equaling the number of unknowns minus two. At 2‐б level, the maximum statistically acceptable value of the MSWD is 1+2(2/f)1/2 (Wendt & Carl 1991). As demonstrated by repeated SL1 standard zircon analyses (weighted mean age of 566.3±2.7 Ma with an MSWD = 0.81, n = 31), our assigned ana‐
lytical random error size is not over‐ or underes‐
timated. Yet, MSWD values for a given sample of our dataset without major external inheritance features are commonly higher than the statisti‐
cally acceptable value unless a significant number of zircons are excluded from the weighted mean age calculation (see below). Following the rea‐
soning of Wendt & Carl (1991) sample MSWD 17
values exceeding its 2‐б range thus should reflect real geological scatter due to external zircon in‐
heritance, antecrystic components, or radiogenic Pb loss; elevated MSWD values may also reflect incorporation of mixed signals ('subtle inheri‐
tance'), i.e., ablating and mixing different age domains of a given zircon, as the dated zircons are commonly small (<100 μm in the longest di‐
mension) compared to applied laser spot diame‐
ters (25‐35 μm). Most of our samples show broad unimodal age peaks in histogram plots and do not allow a clear distinction of multiple popula‐
tions on cumulative probability plots. Thus, zircon age scatter from the sources outlined above in‐
fluences our dataset, but cannot be clearly re‐
solved at our analytical precision; consequently, zircons at the lower and upper age range cannot be unambiguously excluded when calculating the weighted mean age for a given sample. For these reasons, we prefer to include a rather large number of zircon analyses in the calculation of weighted mean ages, unless analyses are clearly rejectable based on histogram distribution criteria or significant age offsets. This practice avoids biasing weighted mean ages by biased zir‐
con age selection aimed at decreasing the MSWD. Weighted mean ages built from a broad zircon population are negligibly susceptible to age bias by radiogenic Pb loss, as these ages of‐
ten show increased individual errors and thus contribute less to the weighted mean age. As it will be shown below, excluding zircons at the lower and upper age range of a given sample of our dataset (where the MSWD of the whole sam‐
ple zircon population exceeds the statistically acceptable value) generally allows driving down the MSWD to a statistically acceptable value without affecting the weighted mean age beyond analytical errors. Weighted mean ages obtained in this manner thus correspond to a mixture of auto‐ and antecrystic age components reflecting the terminal stages of incremental pluton growth. Results Terminology In the following sections we repeatedly use the terms zircon ‘autocryst’, ‘antecryst’ and 18 ‘xenocryst’ which are defined as follows: ‘auto‐
crysts’ refer to zircons grown from the youngest melt batch participating in magma chamber con‐
struction prior to eruption or final pluton solidifi‐
cation (Miller et al. 2007). In contrast, ‘antecrysts’ crystallized from earlier melt batches contribut‐
ing to incremental pluton growth (Miller et al. 2007). The maximum age difference between antecrysts and autocrysts relates to the duration of magmatism at a single volcanic center or plu‐
ton (up to a few m.y.; Miller et al. 2007), or on a regional scale (up to >10 m.y.; Bryan et al. 2008). ‘Xenocrysts’ are incorporated into the magma from genetically unrelated wall rock units (Miller et al. 2007; Bryan et al. 2008). It is important to note that porphyry intrusions (dikes or stocks, as applicable for several samples below) are considered to represent single, rapidly quenched melt batches (e.g., Seedorff et al. 2005). While they thus might contain antecrystic or xenocrystic zircons derived from their parental melts, they are not affected by later melt replen‐
ishment, protracted residual melt crystallization, and incremental pluton growth as large phan‐
eritic intrusions potentially are (e.g., Schaltegger et al. 2009). TIMS analysis A combined TIMS and LA‐MC‐ICP‐MS results summary is presented in Table 3 whereas de‐
tailed TIMS results are shown in Table A1 (Ap‐
pendix). Concordia plots of individual samples, together with weighted mean 206Pb/238U age dia‐
grams are shown in Figure 4. For the following results presentation reported single grain and weighted mean ages are always 206Pb/238U ages where errors include decay constant uncertain‐
ties such that these ages can be compared to ages obtained from other isotopic systems. Due to extremely low radiogenic Pb contents, sample zircons analyzed by TIMS generally show low ra‐
diogenic/common Pb ratios (<1 for most zircons) producing comparatively high age uncertainties on single zircon analyses. As the MSWD of most of our 206Pb/238U ages is significantly below unity, the magnitude of our analytical errors might pos‐
sibly be overestimated. However, MSWD values presented below always overlap with the ex‐
pected MSWD uncertainty range (Wendt & Carl 1991) and are thus statistically acceptable.
19
Zircons of the Gaby and Papa Grande horn‐
blende‐plagioclase porphyry intrusions (samples E05083 and E05090) yield weighted mean ages of 20.26±0.07 Ma (n=5) and 19.89±0.07 Ma (n=6) with MSWD values of 0.10 and 0.11, respectively. Four zircons obtained from a dacitic intra‐caldera dome at Quimsacocha (E06017) yield a weighted mean age of 7.13±0.07 Ma (MWSD = 0.23). All zircon analyses are concordant and overlap within error; we interpret the weighted mean ages to approximate the final emplacement of the respective intrusions. Three zircons of the plagioclase‐hornblende por‐
phyry exposed in underground workings of the La Abundancia Mine at Portovelo (E06112) overlap within error and define a weighted mean age of 24.04±0.06 Ma (MSWD = 0.52). Zircon PO2 dis‐
plays negative discordance; the analysis is offset to the left of the Concordia curve (Fig. 4). Re‐
corded isotopic ratios were stable throughout the whole measurement of this sample and do not indicate any kind of mass spectrometric analytical problem. As the analysis is still concordant within error and the 206Pb/238U age does not seem to be significantly affected by this disturbance we in‐
clude this analysis in the weighted mean age. Four zircons of the Balsapamba pluton (granodio‐
rite; E06140) overlap within error and yield a weighted mean age of 21.46±0.09 Ma (MSWD = 0.23) which we interpret to approximate final emplacement pulse of the intrusion. A single zir‐
con gives an age of 21.13±0.28 Ma possibly re‐
flecting postcrystallization radiogenic Pb loss caused by hydrothermal alteration and is there‐
fore excluded; all analyses are concordant. Seven zircons of a hornblende quartz‐diorite porphyry dike (E06131) intruding the pluton were analyzed yielding a somewhat scattered age distribution (Fig. 4): zircons BA5 and BA6 show ages of 21.39‐
21.41 Ma which is identical to the age of the host rock (E06140); we thus consider them as ante‐
crysts derived from the latter which is in agree‐
ment with resorbed core domains observed in some zircon CL images of this sample (Fig. 3). Four slightly younger zircons from this sample (BA1‐4) would then define a weighted mean age of 21.22±0.17 Ma (MSWD = 0.09). The two slightly older zircons could also be included in a 6‐grain weighted mean age for sample E06131 (dashed line in Fig. 4) yielding a weighted mean 20 age of 21.34±0.11 Ma with a statistically accept‐
able MSWD (MSWD = 0.70), indicating that they could be treated as a single population. Thus, al‐
though there is no statistical justification for dis‐
carding these two zircon analyses as antecrysts from the weighted mean age, we do exclude them based on cross‐cutting field relationships and the high precision age obtained on the host lithology. Consequently, we prefer the first op‐
tion and use the age of 21.22±0.17 Ma as em‐
placement age estimate of the porphyry dike, but note that both ages are identical within error. Finally, a single zircon (BA7) of this sample yields an age of 20.86±0.11 Ma, which we interpret as a postcrystallization radiogenic Pb loss feature. Six zircon analyses of the Apuela batholith at Cuellaje (granodiorite; E06206) overlap within error and yield a weighted mean age of 12.87±0.05 Ma (MSWD = 0.26), interpreted to approximate the final emplacement pulse of the granodiorite pluton. Three of these zircon analy‐
ses discordantly plot to the right of the Concordia curve (Fig. 4) possibly reflecting an imperfect iso‐
topic composition used for common Pb blank correction for this sample. To some extent, this offset might also represent an effect of Pa dis‐
equilibrium not accounted for during data reduc‐
tion (Parrish & Noble 2003). The 206Pb/238U age is not significantly affected, however, and the weighted mean age when excluding these three zircon analyses is identical to the weighted mean age of all six zircons. Individual zircon ages obtained from the Chaucha batholith (granodiorite; E07003) do not overlap within error and show a continuous age distribu‐
tion along the Concordia curve ranging from 15.33±0.06 Ma to 14.84±0.07 Ma (Fig. 4). This age distribution might be produced by mixing variable proportions of antecrystic zircon core domains and autocrystic overgrowth rims. In ad‐
dition (or alternatively), antecrystic components of variable age might be present. Zircon crystalli‐
zation over a time range of c. 0.5 m.y. is consis‐
tent with protracted incremental pluton growth as observed elsewhere (e.g., Schaltegger et al. 2009). A maximum age for the final pulse of plu‐
ton emplacement may be estimated form the two youngest (i.e., closest to autocrystic) zircons which overlap within error and yield a weighted mean age of 14.84±0.07 Ma. Five concordant Figure 4: U‐Pb Concordia diagrams and weighted mean 206Pb/238U age plots of samples dated by TIMS. Error bars and error ellipses correspond to 2‐sigma error ranges. Individual weighted mean ages are presented at 95% confi‐
dence limit with relative and absolute uncertainties as indicated; for additional propagation of the decay constant uncertainty see Table 3. Dashed line in plot for sample E06131 corresponds to weighted mean age if analyses BA5 and BA6 are included; see text for further explanation. All diagrams were generated using the Isoplot v.3.31 Excel macro (Ludwig 2003). 21
Figure 4 (continued) 22 zircons of a granodioritic porphyry dike (E07005) intruding the batholith overlap within error and define a weighted mean age of 9.79±0.03 Ma (MSWD = 1.09) which we interpret to approxi‐
mate the age of porphyry dike emplacement. A single zircon (CH2) plots slightly off the main age cluster (Fig. 4); excluding this zircon from the weighted mean does not have any effect on the age but leads to a decrease of the MSWD from 1.09 to 0.26. Zircon CH1 has an age of 10.26±0.19 Ma, not overlapping within error with the main age cluster (Fig. 4), and is thus interpreted as an‐
tecryst. Four zircons of the El Mozo granodiorite porphyry dike (E07018) overlap within error and define a weighted mean age of 16.04±0.04 Ma (MSWD = 0.53) which we interpret to approximate the age of emplacement of the porphyry dike. Two zir‐
cons dated at 16.16±.0.09 Ma and 16.36±0.06 Ma do not overlap within error with the main age cluster (Fig. 4) and are interpreted as antecrysts. Four concordant zircon analyses of the Junin granodiorite porphyry stock (E07032) overlap within error and yield a weighted mean age of 9.01±0.06 Ma (MSWD = 0.36) which we interpret to approximate the age of porphyry stock em‐
placement. A single zircon (JU1) shows an older age of 9.48±0.20 Ma (Fig. 4) and is considered as an antecryst. LA‐MC‐ICP‐MS analysis Results are presented as weighted mean 206
Pb/238U age plots and corresponding histo‐
grams in Figure 5 and in Table 3; detailed analyti‐
cal results are listed in Table A2 (Appendix). In the following results presentation we only refer to random measurement errors; additional sys‐
tematic errors are considered for weighted mean ages presented in Table 3. Zircons obtained from a biotite‐bearing quartz‐
diorite intrusion at Cangrejos (sample E06066) display a broadly unimodal histogram age distri‐
bution with two minor peaks at the flanks of the major age peak, skewed towards younger ages (Fig. 5). We interpret this pattern as an antecrys‐
tic age component in the two oldest grains, and a group of zircons which suffered radiogenic Pb loss at the younger age range. When excluding the two oldest and three youngest analyses the weighted mean age is 26.0±0.6 Ma (MSWD = 2.9). Further exclusion of the next three youngest grains produces a within error identical weighted mean age of 26.2±0.5 Ma with a statistically ac‐
ceptable MSWD of 1.7. As Pb loss or the ante‐
crystic cut‐off age cannot be clearly analytically resolved in these grains we use 26.0±0.6 Ma as a robust estimate for the final pulse of intrusion emplacement. Sample E07011, a strongly altered Saraguro Group felsite, contains a xenocrystic zircon core at 446 Ma with a 34.7 Ma overgrowth rim; a sec‐
ond zircon has a xenocrystic (or antecrystic) core age of 37.1 Ma vs. a 30.3 Ma tip age. The histo‐
gram age distribution is clearly polymodal (Fig. 5). Excluding the 446 Ma core as xenocryst plus the next six oldest analyses as xeno‐/antecrystic zir‐
cons we obtain a weighted mean age of 30.7±0.5 Ma (MSWD = 4.8) for 22 zircon analyses which we interpret as dating the timing of final magma chamber assembly prior to eruption of the felsite; there might be a small age difference to the de‐
positional age of the tuff. Treating the seven youngest ages as Pb loss results in a weighted mean age of 30.9±0.3 Ma for the remaining 15 zircons, with a statistically acceptable MSWD of 0.83; both ages are identical within error. The histogram plot of sample E07023, a horn‐
blende‐biotite granodiorite intrusion north of Zaruma, shows a bimodal age distribution with peaks around 21 Ma and 29.5 Ma (Fig. 5). The latter comprises the five oldest analyses and seems to represent an end member xeno‐ or an‐
tecrystic component yielding a weighted mean age of 29.5±1.0 Ma. The 15 youngest zircon ages define a weighted mean age of 20.7±0.8 Ma with a statistically acceptable MSWD of 1.07 which we use as a proxy for the last pulse of intrusion em‐
placement. Five analyses are transitional be‐
tween the two age groups and might reflect vari‐
able proportions of the older and younger age groups, and/or antecrystic zircon components of variable age. Sample E07030, a Curiplaya hornblende quartz‐
diorite porphyry stock, shows two zircon analyses clearly offset from the bulk of the zircon analyses (Fig. 5). We interpret these two ages as strongly influenced by radiogenic Pb loss and exclude them from further discussion. Otherwise, the 23
24 histogram age distribution is unimodal, albeit slightly skewed towards younger ages suggesting subtle radiogenic Pb loss for some younger grains. The bulk of sample zircons (n=26) gives a weighted mean age of 92.0±1.0 Ma (MSWD = 3.3). Progressively narrowing down the range of analyses by excluding the three oldest and eight youngest ages produces a within error identical weighted mean age of 92.4±0.7 Ma with a statis‐
tically acceptable MSWD of 1.3. We prefer to use the first value of 92.0±1.0 Ma which we use as a proxy for emplacement of the porphyry intrusion. The zircon age histogram of sample E07045, a hornblende‐bearing granodiorite intrusion of the Telimbela‐Chazo Juan pluton, shows a broadly unimodal age distribution which is slightly skewed towards younger ages, possibly as a re‐
sult of radiogenic Pb loss (Fig. 5). Where meas‐
ured separately, zircon core and tip ages overlap within error. Excluding the four oldest grains of this sample, the 24 remaining zircons yield a weighted mean age of 25.5±0.6 Ma with an MSWD of 2.2, where the maximum statistically acceptable MSWD is 1.6; we interpret this age to approximate final intrusion emplacement. Reject‐
ing the five youngest and three oldest zircons of this group yields a within error identical weighted mean age of 25.5±0.5 Ma with a statistically ac‐
ceptable MSWD of 1.2 for the remaining 16 zir‐
cons. Discussion Causes for limited external zircon in‐
heritance in Tertiary intrusions Zircon xenocrysts can become incorporated into arc magma either in its source region, in which case they are referred to as "inherited" (we addi‐
tionally use the qualifier "external" to stress the xenocrystic character of these zircons, and to set them apart from recycled antecrysts related to the same magmatic system), as well as during magma ascent and later‐stage mid‐shallow crustal differentiation. The presence of inherited xenocrystic zircon cores with thick magmatic overgrowth rims is usually indicative of constant zircon‐melt immersion and zircon saturation of melts from the magma source region onwards. Due to rapid dissolution kinetics, preservation of inherited zircons to significant amounts in zircon‐
undersaturated melts is otherwise only possible if these zircons become encapsulated in other crys‐
tallizing minerals (e.g., Hansmann & Oberli 1991; Miller et al. 2003). In a study of a wide range of granitoid intrusions in different geodynamic envi‐
ronments, Miller et al. (2003) observe a bimodal distribution pattern with either external zircon inheritance‐poor or ‐rich granitoids; the former are mainly of calc‐alkaline type, mostly meta‐
aluminous, and source melt temperatures are on average at least 837°C. Our TIMS and LA‐MC‐ICP‐MS age data suggest that the occurrence of externally inherited zircon cores in all intrusions investigated in this study is extremely limited. This is in marked contrast to some Paleozoic intrusions in Ecuador where ex‐
ternal zircon inheritance is an abundant feature (Noble et al. 1997). Zircon textural analysis by SEM‐CL imaging prior to TIMS analysis demon‐
strates the relatively abundant occurrence of dis‐
tinct zircon core domains in some Tertiary intru‐
sions (Fig. 3; Tab. 2). These domains do not rep‐
resent xenocrystic, externally inherited cores, but instead seem to reflect antecrystic zircon, or Figure 5 (previous page): Weighted mean 206Pb/238U age plots, histograms, and cumulative probability density function
curves of samples and SL‐1 standard zircon analyzed by LA‐MC‐ICP‐MS. Effects of excluding certain groups of zircons from the calculation of the weighted mean age are illustrated in the plot (outlined with red lines) and discussed further in the text. Preferred weighted mean ages are marked by thin black horizontal lines. Empty boxes = rejected zircon analyses based on histogram age distribution (xenocrysts, Pb loss); gray boxes = zircon analyses potentially influenced
by analytically not clearly resolvable antecrystic components or Pb loss; black boxes = zircon analyses accepted for weighted mean age calculation. Weighted mean ages displayed in bold, calculated from zircon analyses marked in gray
and black, are preferred as proxies for the final phase of intrusion emplacement (or magma chamber assembly prior to
eruption); their MSWD values are not always statistically acceptable for a single zircon population (Wendt & Carl 1991).
Weighted mean ages not displayed in bold are calculated only from zircon analyses marked in black and always corre‐
spond to statistically acceptable MSWD values. Note that both weighted mean ages for a given sample are generally identical within error. Also shown are Concordia plot and weighted mean age of SL‐1 standard zircon; the MSWD of 0.81 for the SL‐1 weighted mean 206Pb/238U age suggests analytical random errors are estimated accurately. Error bars,
error ellipses, and weighted mean uncertainties correspond to analytical random errors at 2‐ level; see Table 3 for propagation of additional systematic errors. All diagrams generated using the Isoplot v.3.31 Excel macro (Ludwig 2003).
25
distinct autocrystic zircon growth sequences which do not show significantly different isotopic age domains, and do not change U‐Pb age con‐
cordancy between the 238U/206Pb and 235U/207Pb isotopic systems. Zircon core and tip ages ob‐
tained by LA‐MC‐ICP‐MS are mostly indistin‐
guishable within error. Zircon selection for TIMS analysis in this study was restricted to picking fresh‐looking, inclusion‐free zircons to minimize alteration‐induced Pb loss and high common Pb contents. Single‐grain hand picking for mass spec‐
trometric analysis did not discriminate zircons based on their morphologies, such that a rela‐
tively representative cut of the sample zircon in‐
ventory was obtained. Thus, the observed deficit of externally inherited zircon components sug‐
gested by TIMS age data does not seem to be a function of morphology‐selective zircon hand‐
picking, and is mirrored by our LA‐MC‐ICP‐MS results. The extremely limited occurrence of ex‐
ternal zircon inheritance in Tertiary intrusions might be due to the following reasons: (1) Oceanic plateau basement units (or island arc intrusive roots) of the Western Cordillera and the Interandean region of Ecuador are dominantly primitive and thus zircon‐poor (e.g., Spikings et al. 2005; Chiaradia et al. 2009; Vallejo et al. 2009). For the most part, detrital zircon‐bearing sedimentary formations partly sourced from landward metamorphic regions constitute the only potentially effective assimilant source for zircon xenocrysts in the Western Cordillera, as has been demonstrated for the Yunguilla and Sa‐
guangal formations, as well as for the sedimen‐
tary portions of the Macuchi Unit (Vallejo 2007). A Late Cretaceous Curiplaya porphyry intrusion, situated in the Celica‐Lancones basin north of the Tangula batholith (Fig. 1), does not contain any inherited zircon component indicating a domi‐
nantly primitive crustal basement, as opposed to the mature continental basement in the El Oro range further north (see also Chapter 4). (2) Ascending melts repeatedly exploited the same structures through time (Tab. 4, and further discussion in the Appendix). Crystallization of arc intrusive root zones along these structures shielded ascending and accumulating melt batches from contamination by xenocrystic zir‐
26 cons such that (if zircons were present; see above) assimilation of zircon antecrysts (sensu Bryan et al. 2008) originating from arc intrusive roots was favored. This scenario is in agreement with the relatively common occurrence of ante‐
crystic zircon domains in several samples investi‐
gated in this study, and is further compatible with the reasoning of Dungan & Davidson (2004) where crustal contamination of arc magmas in a broadly stationary magmatic arc complex tends to be restricted to arc intrusive root zones of similar isotopic composition. (3) Hot, mantle‐derived melts are often zircon‐
undersaturated (Miller et al. 2003). In the case of Ecuador, this could apply to deep‐mid crustal domains where hot, mantle‐derived melts tem‐
porally stall and evolve, i.e., the MASH zone of Hildreth & Moorbath (1988), or the hot zone of Annen et al. (2006). Granitoid parental melts would only become zircon‐saturated with de‐
creasing temperature shortly prior to or during their final emplacement in the shallow crust. In a single sample dated in this study, a felsite at Tres Chorreras assigned to the Saraguro Group (E07011; 30.7±1.6 Ma), a few Early Oligocene xeno‐ or antecrystic zircons are present, and a single Ordovician xenocrystic zircon core age with a thick Early Oligocene overgrowth rim was de‐
tected. The latter indicates prolonged melt expo‐
sure of the xenocrystic core, suggesting the melt was zircon‐saturated. As only a single core of this age was identified, it likely does not represent external source inheritance but rather contami‐
nation during magma ascent. Vallejo (2007) stud‐
ied detrital zircons of the Yunguilla Formation and identified five populations spanning a total age range from the Mesozoic to the Precambrian. His subpopulation B1 (384‐639 Ma) overlaps in age with the inherited Ordovician core at Tres Chorreras. Given that the Cretaceous Yunguilla turbidites were deposited in a wide paleo‐forearc basin corresponding to the present‐day Western Cordillera and Interandean region (Vallejo 2007), this formation could likely constitute an assimi‐
lant for the Tres Chorreras (Saraguro) parental melt. Geodynamic controls on Tertiary arc magmatism In order to evaluate feedback reactions between arc magmatism and the geodynamic regime at the Tertiary Ecuadorian margin, we compiled K‐Ar and ZFT data from a large number of sources for Tertiary volcanic and plutonic rocks in Ecuador, in addition to U‐Pb data presented in this study (Tab. A3 in the Appendix). As discussed in the Appendix, K‐Ar data can largely serve as an accu‐
rate proxy for arc magmatism on a regional, multi‐m.y. scale, although at individual magmatic centers disturbed K‐Ar ages may occasionally oc‐
cur (e.g., at Apuela/Junin). Where K‐Ar and U‐Pb data for the same lithology are available, maxi‐
mum age differences between the two methods are on the order of 1‐4 m.y. Potentially inaccu‐
rate and disturbed ages were omitted from the data base (Tab. A3) such that the data used for the following discussion are thought to be accu‐
rate for the timing of magmatism on a regional scale. Tertiary convergence rates and obliquities at the Ecuadorian margin We computed Mid‐ to Late Tertiary (40‐0 Ma) convergence velocities and obliquities at 2°S/82°W, corresponding to the Central Ecuador‐
ian trench (Fig. 6), using the most recent set of available Farallon/Nazca‐South America rota‐
tional parameters (Somoza 1998) and the UNAVCO online plate motion calculator (http://sps.unavco.org/crustal_motion/dxdt/mod
el). Along‐arc variations in convergence velocity and obliquity due to the age‐specific position of a given rotation pole are negligible for the small latitudinal difference between southern‐ and northernmost Ecuador (5°S to 1°N). Convergence parameters in Figure 6 are plotted without uncer‐
tainty ranges because Somoza (1998) does not quantify the uncertainties associated with his rotation poles. Pardo‐Casas & Molnar (1987) pre‐
sent relative convergence rate uncertainties of c. 25% for Miocene Farallon/Nazca‐South America convergence rates around 100 mm/y at 10°S. Since Somoza’s (1998) study is based on the same plate reconstruction systematics as the approach of Pardo‐Casas & Molnar (1987) it seems justified to apply identical relative uncertainties as a first‐
order estimate, although Somoza (1998) notes that he expects his rotational parameters to be slightly more precise, mainly due to improved accuracy and resolution of the geologic time‐
scale. Assuming 25% relative uncertainties for conver‐
gence rates plotted in Figure 6 results in within error overlapping rates throughout most of the Miocene, but allows identification of a peak of nearly orthogonal plate convergence at >120 mm/y in the Late Oligocene‐Early Miocene. Com‐
pared to Late Eocene‐Early Oligocene conver‐
gence rates, the significant acceleration and obliquity change starting at 28.3‐25.8 Ma and culminating in the 25.8‐20.2 Ma period is in agreement with the postulated change in Faral‐
lon plate motion prior to its fission at 24 Ma and initiation of Cocos‐Nazca seafloor spreading at 23 Ma (Lonsdale 2005; Barckhausen et al. 2008). The change in both convergence rate and obliquity in the 20.2‐16.0 Ma period cannot be clearly re‐
solved at 25% relative uncertainty. It might relate to a short‐term variation caused by the estab‐
lishment of post‐fission independent Nazca plate motion commencing at c. 20 Ma when the east‐
ward‐propagating plate rupture eventually inter‐
sected the Meso‐South American trench system; Figure 6: Convergence parameters from 40‐0 Ma at the Ecuadorian margin. Calculated at 2°S/82°W (present‐
day central Ecuadorian trench position) using the Faral‐
lon/Nazca‐South America rotational poles of Somoza (1998). Note discussion of associated uncertainties in the text. Also shown are major tectonic events affect‐
ing the Farallon/Nazca plate during the plate reorgani‐
zation in the Oligocene‐Miocene (Lonsdale 2005; Barckhausen et al. 2008) which comprise: 1 ‐ Farallon plate fission initiates at Farallon‐Pacific spreading cen‐
ter and propagates towards South America; 2 – Cocos‐
Nazca seafloor spreading initiates; 3 – plate rupture intersects South American trench; independent Nazca and Cocos plate motion starts. 27
Figure 7: Pluton and volcanic radiometric age versus latitude plot of southern Ecuador (SE of the CPPF) illustrating along‐arc migration patterns of arc magmatism. Screened radiometric age references are shown in Table A3 (Appen‐
dix); geologically inaccurate ages and duplicate samples were omitted from the database. Minimum ages progressively young northwards (red arrow) from the Peruvian border region (Portachuela batholith) towards the southern end of
the present‐day Northern Volcanic Zone (Sangay volcano), likely reflecting northwards slab flattening due to the c. 14‐
10 Ma inception of the Inca plateau at the northern Peruvian/southern Ecuadorian margin (Gutscher et al. 1999a;
Rosenbaum et al. 2005) following a period of protracted arc magmatism throughout most of the Tertiary. this process was accompanied by a reorientation of the young Cocos‐Nazca spreading center (Barckhausen et al. 2008). Along‐arc distribution of Tertiary arc magmatism The Tertiary latitudinal migration pattern of arc magmatism through time for southern Ecuador (SE of the CPPF) is shown in Figure 7. Minimum ages of exposed volcanic and plutonic rocks pro‐
gressively young northwards. The youngest phases of active arc magmatism in southernmost Ecuador are recorded by the Portachuela batho‐
lith (c. 4°30'‐5°S) and the Tarqui Formation (c. 3°30'S) with ages around 10‐12 Ma; younger Tar‐
qui units occur only further north (e.g., Hunger‐
bühler et al. 2002). The cessation of active arc magmatism in northern Peru and southern Ecua‐
dor is temporally and spatially associated with the establishment of a flat slab subduction set‐
ting due to subduction of the buoyant Inca pla‐
teau starting at 14‐10 Ma (Gutscher et al. 1999a; Rosenbaum et al. 2005). The northward decrease in minimum ages of arc magmatism towards the southernmost active arc volcano in Ecuador, San‐
gay, thus probably reflects the progressive northward broadening of the flat slab region (see Fig. 2 for present‐day extent). A single ZFT age of 28 2.3±0.8 Ma for pyroclastic rocks of the Salapa Formation near Loja in southern Ecuador (Hun‐
gerbühler et al. 2002) could be in conflict with this model if accurately dating a magmatic event. The mixed average ZFT age of this sample is 16.4±7.4 Ma (n = 23) with a χ2 probability of 0%, leading these authors to infer a depositional age of 2.3±0.8 Ma based on the youngest zircon. Given these statistical ambiguities and the in‐
tense alteration features displayed by the Salapa volcanics (Hungerbühler et al. 2002) the inferred Salapa age might by inaccurate and we therefore exclude it from our database. Across‐arc distribution of Tertiary arc magma‐
tism Bulk across‐arc migration patterns may be used to track changes in slab dip, although structurally controlled magma ascent and emplacement, as well as crustal thickening additionally influence arc magmatic outcrop patterns (e.g., Trumbull et al. 2006). Figure 8 illustrates Tertiary longitudinal migration patterns of arc magmatism through time, combined with radiometric age histograms (U‐Pb, K‐Ar, and ZFT data from screened data‐
base; see discussion below) for two major Ecua‐
dorian arc segments. Data for southern‐ and northernmost Ecuador (4‐5°S and 1°N‐1°S, re‐
spectively) are not shown in Figure 8 as for the former the scarcity of available data does not allow a reliable analysis of across‐arc migration patterns, and for the latter the present‐day trench obliquity is high and partly irregular, the paleo‐trench configuration cannot be predicted with confidence, and, consequently, the age‐
longitude distribution of this arc segment is not directly comparable to the others. relate to slab steepening (increasing hot astheno‐
sphere convection into the mantle wedge, and, ultimately, increased crustal melting; see discus‐
sion below), similarly to what is inferred in parts of the central and southern Andes at that time (e.g., Mamani et al. 2010). However, the rear arc position in all Ecuadorian arc segments remains relatively stable (Fig. 8) until the Late Miocene instead of systematically migrating westwards as might be expected from slab steepening. The radiometric age base in northern‐central Ec‐
uador (100 km arc strike length) is dominated by plutonic rock samples. Regional Andean‐ (NNE) trending structures (mainly the CTSZ) exert a ma‐
jor control on pluton emplacement in Ecuador (Fig. 2 and Tab. 4; further discussion in the Ap‐
pendix) such that the relative stability of the Oli‐
gocene‐Miocene across‐arc position (Fig. 8) may partly relate to structurally controlled pluton em‐
placement, and the overall arc magmatic outcrop pattern may be biased trenchwards as major structures dip 35°E at deep to mid‐crustal levels (Guillier et al. 2001). Figure 8 shows that signifi‐
cant eastward frontal arc migration occurred in the Mid‐ to Late Miocene; major rear arc east‐
ward broadening is recorded for the Late Mio‐
cene‐Pliocene. Eastward arc migration and broadening during the Late Miocene‐Pliocene in all Ecuadorian arc segments is consistent with minor‐moderate slab flattening, although eastward arc migration (but not broadening) could in addition partly relate to subduction erosion which affects the present‐day Ecuadorian margin (e.g., Sage et al. 2006). The timing of the landward arc front migration in the Late Miocene can be traced by the youngest plu‐
tons exposed in deeply eroded parts of the Western Cordillera: c. 6 Ma marks the youngest magmatism in northern Ecuador (Junin porphyry intrusions), whereas the youngest major plutonic activity in central Ecuador occurred at c. 14‐15 Ma (Telimbela‐Chazo Juan and Corazon intru‐
sions), either indicating Late Miocene‐Pliocene eastward arc migration might have proceeded slightly diachronously along the arc, and/or post‐
14 Ma plutons exist, but have not been dated yet in the Western Cordillera of central Ecuador. The age of landward arc migration in northern‐central Ecuador broadly coincides in time with the arrival of the Carnegie Ridge seamount chain at the Ec‐
uadorian‐Colombian trench at c. 8 Ma (Daly 1989; Gutscher et al. 1999b; Chapter 2); the latter might induce moderate slab shallowing by virtue of its buoyancy (e.g., van Hunen et al. 2004). Arc migration patterns do not indicate any significant pulses of pre‐Late Miocene subduction erosion affecting the Ecuadorian margin. The radiometric age database for southern‐
central Ecuador (200 km arc strike length) com‐
prises a large number of widespread volcanic and plutonic samples in subequal proportions. Data for this arc segment are thus potentially more sensitive to record arc migration patterns than data for northern‐central Ecuador. The age‐
longitude distribution of plutons and volcanics in southern‐central Ecuador suggests c. 50 km of arc broadening by a westward frontal arc jump in the Late Oligocene while the rear arc position re‐
mained relatively fixed (Fig. 8). Arc broadening is reversed by progressive (?) eastward frontal arc migration in the Mid‐Miocene and culminates in a Late Miocene‐Pliocene period of eastward arc migration comprising both the frontal and rear arc regions, and mirroring the Late Miocene age‐
longitude distribution trend in northern‐central Ecuador. Although partly structurally controlled, major westward arc broadening in southern‐central Ec‐
uador in the Late Oligocene (Fig. 8) might thus Productivity of Tertiary arc magmatism Radiometric age histograms (Fig. 8) have been used as proxies for plutonic and volcanic rock volumes (e.g., Glazner 1991). They can only serve as first‐order proxies for arc magma production because they reflect arc exposure and erosion conditions as well. Both major Ecuadorian arc segments display age peaks initiating in the Late Oligocene‐Early Mio‐
29
cene (yellow boxes in Fig. 8). Only in the northern part of the northern Ecuadorian arc segment (not shown), the peak initiates slightly later in the Early to Mid‐Miocene. The peak in volcanism es‐
sentially corresponds to the widespread eruption of upper Saraguro Group volcanics which mainly comprise ignimbrites in its upper portion (Tab. 1). The peak in plutonism dominantly reflects the construction of the central Ecuadorian batholith system, the Cangrejos‐Zaruma intrusive belt, and associated smaller intrusions. Combined, this peak distribution suggests a relative arc mag‐
matic flare‐up event in Ecuador during the Late Oligocene‐Early Miocene. It is important to note that, as we do not attempted to quantify arc magma production rates at these times, this flare‐up event may not correspond to the scale of an arc magmatic flare‐up event in the sense of, for example, Ducea & Barton (2007) with magma production rates > 75‐100 km3/m.y. arc km‐1. Here, we are using the term flare‐up to indicate qualitatively a significant increase in arc magma‐
tism compared to the Eocene‐Early Oligocene. A second peak in volcanism occurs in the Late Mio‐
cene and probably reflects the increasing preser‐
vation potential of young volcanic sequences, and eastward arc broadening associated with moder‐
ate slab flattening. Figure 8: Pluton and volcanic radiometric age versus longitude plots and age histograms illustrating the Tertiary arc position (migration marked by red arrows) and arc magmatic intensity in Ecuador (from 1‐4°S). Used radiometric ages (U‐Pb, K‐Ar, ZFT) were carefully screened, and potentially inaccurate or duplicate ages have been removed from the
database (Appendix Table A3). The Paleocene‐Eocene arc position is relatively stable in all arc segments. Arc broaden‐
ing occurs by a c. 50 km westwards jump of the frontal arc during the Late Oligocene in southern‐central Ecuador; this is not clearly observed in the arc segment further north because of structurally controlled pluton emplacement (along the CTSZ and CPPF) and almost complete erosion of Oligocene‐Miocene volcanic rocks. Progressive eastward migration of the frontal arc at constant rear arc positions during the Early to Mid‐Miocene is followed by a Late Miocene‐Pliocene period of significant eastward broadening of the rear arc. Age histograms show Late Oligocene‐Early Miocene peaks in arc magmatism in all arc segments indicative of a regional arc magmatic flare‐up event (yellow boxes). 30 Glazner (1991) suggests that periods of intense plutonism might be associated with oblique con‐
vergence settings as the latter result in crustal strike‐slip deformation and thus space creation for pluton emplacement. We do not observe any systematic correlations between convergence obliquity (Fig. 6) and plutonism (Fig. 8) in Tertiary Ecuadorian arc magmatism. However, as the Ec‐
uadorian margin represents an oblique subduc‐
tion system (where obliquity is variably accom‐
modated by oblique subduction slip and crustal strain partitioning; Ego et al. 1996 and Appendix) the Tertiary margin has probably undergone large‐scale crustal strike‐slip deformation throughout the Tertiary such that plutonism was principally continuous. Peak periods of pluton emplacement might in part, however, relate to reactivation of large strike‐slip fault systems. Asthenospheric controls on a Late Oligocene‐
Early Miocene flare‐up event in Ecuadorian arc magmatism Mamani et al. (2010) attribute widespread Late Oligocene‐Early Miocene ignimbrite eruptions in the central Andes to increased crustal melting and asthenospheric melt production in response to slab steepening. As discussed above, Late Oli‐
gocene‐Early Miocene arc broadening in south‐
ern‐central Ecuador, and the correlated peaks in arc volcanism and plutonism might also be asso‐
ciated with slab steepening, although a system‐
atic migration of the rear‐arc position is not ob‐
served. In addition, the flare‐up event in Ecuador coin‐
cides with a major acceleration in Farallon/Nazca‐
South America convergence rates at the Ecuador‐
ian trench (Fig. 6). This suggests a feedback mechanism between the arc magmatic flare‐up event and plate tectonics operates at an astheno‐
spheric scale. The same has been proposed for the southern Chilean margin for the same time interval during which neither subducting slab properties nor overriding plate motion under‐
went any major changes (Jordan et al. 2001). Melt production in the mantle wedge is concen‐
trated in the region where slab dehydration‐
derived volatiles first encounter fertile mantle material of a sufficiently high temperature to in‐
duce partial melting (e.g., Cagnioncle et al. 2007; note that, additionally, decompression melting might take place in other regions). Subsequently, the melt fraction may be modified by interaction of the ascending melt with the surrounding man‐
tle peridotite (e.g., Grove et al. 2003). Several parameters control increased mantle melting and, by inference, increased arc magma produc‐
tion and development of a broader arc at faster convergence rates and/or during slab steepening. These essentially comprise variations in slab‐
derived volatile flux, the volatile fraction reaching the zone of partial melting in the supra‐slab Table 4: Regional structures in Ecuador associated with Tertiary intrusions investigated or
referenced in this study
Structure
Chimbo-Toachi
shear zone
(CTSZ)
associated intrusions
Santiago, Apuela-Nanegal
(-Junin/Cuellaje), Corazon,
Telimbela-Chazo Juan,
Balsapamba-Las Guardias,
Echeandia
remarks
extends to deep to mid-crustal levels (c. 35°E dip; Guillier et al.
2001); originally regarded as suture zone for Macuchi island arc
(Hughes & Pilatasig 2002), but more recently dismissed as suture
with autochtonous Macuchi origin (Vallejo 2007; Vallejo et al. 2009)
Calacalí-PujiliPallatanga fault
zone (CPPF)
Chaucha
extends to deep to mid-crustal levels (c. 35°E dip; Guillier et al.
2001); western limit of regional suture zone between accreted oceanic plateau units and the paleocontinental margin (Spikings et al.
2005; Vallejo et al. 2009)
northern Amotape suture zone
Cangrejos-Zaruma intrusive
belt
bracketed between Piñas-Portovelo and Jubones fault systems;
Piñas-Portovelo fault joins westwards with La Palma-El Guayabo
and Tahuin Dam (Naranjos) faults, delimiting the deeply exhumed
metamorphic Raspas complex whose structural position has been
related to the ancient Amotape suture zone (Bosch et al. 2002)
See Appendix for a more detailed discussion on structurally controlled intrusion emplacement.
31
mantle wedge by porous flow, mantle wedge temperatures and the (lateral) extent of the zone where wet melting occurs, as well as the supply rate of fertile mantle material from induced re‐
turn flow (Fig. 9). The relative importance of each parameter may vary depending on which bound‐
ary conditions apply (Cagnioncle et al. 2007). In a simplified two‐dimensional numerical mantle flow model of a slice oriented parallel to the sub‐
duction slip, Kincaid & Sacks (1997) show that a temperature‐controlled viscous boundary layer forms in the asthenosphere adjacent to the downgoing slab, where mantle material is dragged downwards parallel to the slab at rates proportional to the subduction slip. Removal of asthenospheric boundary layer material from the mantle wedge is balanced by forced return flow of hot, fertile mantle material sourced from a lower backarc region (Kincaid & Sacks 1997); in a more realistic three‐dimensional environment return flow sources could additionally be dy‐
namically distributed in along‐arc dimension (e.g., Behn et al. 2007). The return flow rate, coupled to the slab velocity, thus directly controls two main parameters for partial melt production in the mantle wedge, namely the supply of fertile mantle material and mantle wedge temperatures (Cagnioncle et al. 2007). Kincaid & Sacks (1997) demonstrate that the maximum mantle wedge temperature increases as a function of subduction slip; the effect is not linear, but becomes more significant for changes Figure 9: Simplified schematic subduction zone cross section of the Ecuadorian margin as broadly applicable for Late Oligocene‐recent times illustrating multiple stages of arc magma petrogenesis. Progressive slab dehydration takes place below the forearc and main arc regions. Temperature‐controlled formation of a viscous boundary layer couples asthenospheric material with slab motion, and induces return flow into the mantle wedge. In simplified two‐
dimensional models, melt production in the mantle wedge is controlled by the amount of slab‐derived fluids reaching the region of partial melting in the mantle wedge, and by asthenospheric return flow rates providing fertile mantle ma‐
terial and possibly increasing mantle wedge temperatures (e.g., Cagnioncle et al. 2007). These processes operate more vigorously at higher subduction slip velocities and may thus increase partial melt production in the mantle wedge. Note
that in a three‐dimensional environment additional along‐arc controls may influence the productivity of partial melting
and arc magmatism (e.g., Tamura et al. 2002). Major fault geometries shown in the Figure correspond to the present‐
day configuration based on Guillier et al. (2001); fault systems of the sub‐Andean zone and the Eastern Cordillera are not shown. Upper plate Tertiary pluton distribution suggests strong structural control of major structures (partly su‐
tures) on magma ascent (possibly in part non‐vertical) and pluton emplacement (see Tab. 4 and further discussion in
the Appendix). General petrogenetic aspects of cross section adapted from Kincaid & Sacks (1997), Stern (2002), Grove et al. (2003), Annen et al. (2006), and Cagnioncle et al. (2007). 32 from slow‐moderate to fast subduction slips, whereas it becomes nearly negligible for velocity increases taking place at already high subduction slips (> c. 100mm/y, depending on the overriding plate thickness). Calculating the due‐east Faral‐
lon/Nazca‐South America convergence rates (from Fig. 6) as a first‐order proxy for subduction slip at the Ecuadorian margin reveals a Late Oli‐
gocene‐Early Miocene increase from 60‐70 mm/y to 120‐140 mm/y thus representing a critical in‐
crease from moderate to high subduction slips where a significant effect on maximum mantle wedge temperature would be expected. Higher plate convergence rates (and more trench‐orthogonal convergence), if proportional to local subduction slip velocities, might increase the amount of volatiles introduced into the man‐
tle wedge by slab dehydration, but at the same time decrease the volatile fraction reaching the region of partial melting by increased downwards volatile advection due to more vigorous return flow; the latter effect becomes less pronounced at higher mantle wedge permeability, i.e., faster volatile migration rates (Cagnioncle et al. 2007; see also Zellmer 2008). In two‐dimensional mod‐
els, faster convergence rates result in higher rates of mantle partial melt production where the proportionality between the two may vary as a function of the interplay of various control pa‐
rameters (Cagnioncle et al. 2007). Consequently, the arc magmatic flare‐up event observed in Ec‐
uador during the Late Oligocene‐Early Miocene might ultimately reflect increased asthenospheric melt input into the arc crust driven by increased plate convergence rates, possibly further accen‐
tuated by slab steepening. With increasing return flow rates, mantle wedge partial melts may be advected closer towards the trench, thus princi‐
pally causing trenchward arc broadening as ob‐
served in the Saraguro arc segment during the Late Oligocene‐Early Miocene; however, the vol‐
canic front position might also be controlled by additional factors such as melt collection in a de‐
compaction channel (Cagnioncle et al. 2007). It is important to note that along‐arc heterogeneities in mantle wedge partial melt production are likely to further influence arc magmatic distribu‐
tion patterns (e.g., Tamura et al. 2002); the latter are not considered in the simplified two‐
dimensional models discussed above. Crust‐mantle wedge feedback impacts on arc magmatism Increased partial melt production in the mantle wedge heats the crust of the overriding plate by advection (due to a larger volume of ascending partial melts) and by increased conductive basal heat flow into the overlying South American lithosphere. This implies a higher melt fraction in a deep crustal hot zone (Annen et al. 2006) and could increase the crustal contribution to arc magmatism by lowering the relative thermal threshold for assimilation. Increased crustal melt‐
ing is inferred for arc magmatic flare‐up events elsewhere, and might also be driven by additional factors such as crustal thickening or lithospheric delamination (e.g., Ducea & Barton 2007). The widespread occurrence of Late Oligocene‐Early Miocene silicic ignimbrites as part of the upper Saraguro Group (Tab. 1) is consistent with in‐
creased crustal melting. Whole‐rock isotopic compositions of Tertiary arc units do not show any straightforward systematic variations in the Late Oligocene‐Early Miocene indicative of bulk increased continental crust contributions (Chap‐
ter 4). However, this is anticipated as mid‐ to deep crustal basement units of the Tertiary arc in Ecuador are mostly composed of isotopically primitive oceanic material, and spatio‐temporal isotopic variations mainly relate to the tectoni‐
cally controlled occurrence of continental base‐
ment units, and the vertical level of arc magma differentiation in the crust, instead of the total amount of crustal contamination (Chapter 4). Increased heat flow into the overriding plate trig‐
gered a regional uplift event and resulted in re‐
gional horizontal extension along the Chilean main arc in the Late Oligocene‐Early Miocene (Jordan et al. 2001). This tectonic setting is con‐
sistent with the inferred tensional environment during deposition of the Saraguro Group (e.g., Steinmann 1997) and correlates in time with widespread elevated cooling and exhumation rates along the Ecuadorian margin in the Mio‐
cene, as inferred from thermochronologic model‐
ing (e.g., Spikings et al. 2005). 33
Conclusions This study presents the first dataset of robust U‐
Pb zircon ages on Late Tertiary intrusive rocks of Ecuador. The regional distribution trends of Ter‐
tiary plutons at the Ecuadorian margin mirror the along‐ and across‐arc orientations of deeply‐
rooted major fault zones suggesting crustal magma ascent and intrusion emplacement were principally controlled by these structures, further modulated by distributed (mainly transpres‐
sional) shear in the upper crust. External zircon inheritance in Tertiary intrusions of the Western Cordillera and the Interandean region of Ecuador is very minor, likely reflecting dominantly zircon‐
poor oceanic basement units as potential assimi‐
lants. In addition, zircon assimilation during magma ascent and mid‐ to shallow crustal differ‐
entiation was mainly restricted to antecrysts de‐
rived from intrusive root zones, likely as a result of continuous preferential channeling of arc magmas through the same, deeply‐rooted struc‐
tural conduits. Where both K‐Ar and U‐Pb data exist for a given lithology, ages obtained by the different methods are usually concordant within 1‐4 m.y. implying that K‐Ar data may be used as a semi‐accurate proxy for the timing of Tertiary arc magmatism in Ecuador on a broad, regional scale. Combining previously published K‐Ar and ZFT age informa‐
tion with newly obtained zircon U‐Pb ages of Ter‐
tiary intrusions allows us to identify a pro‐
nounced pluton emplacement peak in the Late Oligocene to Early (Mid‐) Miocene. The peak in plutonism temporally coincides with the aerially extensive eruption of Saraguro Group ignim‐
brites, and seems to indicate a regional, transient arc magmatic flare‐up event in Ecuador involving significant westward arc broadening in central Ecuador. At the time scale resolvable by plate tectonic reconstructions, initiation of the arc magmatic flare‐up event coincides with a change in Farallon plate motion prior to its fission involv‐
ing a significant acceleration of Farallon/Nazca‐
South America convergence rates, and suggesting a positive feedback operating between faster plate convergence rates, asthenospheric melt production and arc magmatic productivity in Ec‐
uador. This might be envisaged by a higher amount of slab‐derived fluids introduced into a 34 given mantle wedge volume, and by a change in asthenospheric flow dynamics where widespread positive thermal anomalies develop in the subarc mantle wedge and replenishment rates of fertile mantle material increase in response to more vigorous induced return flow. The latter might additionally be influenced by slab steepening. Therefore, we attribute the flare‐up event in arc magmatism to increased melt production in the mantle wedge causing increased mantle‐crust melt flux and increased heat transfer into the crust of the overriding plate. As elsewhere (e.g., Ducea & Barton 2007) a positive tectonomag‐
matic‐thermal feedback mechanism may be in‐
duced, facilitating increased partial crustal melt‐
ing, voluminous magma storage at upper crustal levels leading to batholith construction, and/or ignimbrite eruption at the Earth’s surface. How‐
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tions. In: Oncken, O. et al. (eds.), The Andes, Springer, p. 29–43. Vallejo, C. (2007): Evolution of the Western Cordillera in the Andes of Ecuador (Late Cretaceous‐Paleogene). PhD Thesis, ETHZ, Zürich, Switzerland, 208 p. Vallejo, C., Winkler, W., Spikings, R. A., Luzieux, L., Heller, F., Bussy, F. (2009): Mode and timing of terrane accretion in the forearc of the Andes in Ecuador. In: Kay, S.M., Ramos, V.A., Dickinson, W. R. (eds), Back‐
bone of the Americas: Shallow Subduction, Plateau Uplift, and Ridge and Terrane Collision. Geological Society of America Memoir 204. Van Hunen, J., van den Berg, A. P., Vlaar, N. J. (2004): Various mechanisms to induce pre‐sent‐day shallow flat subduction and implications for the younger Earth: a numerical parame‐ter study. Physics of the Earth and Platenary Interiors 146; 179‐194. Van Thournout, F. (1991): Stratigraphy, magmatism and tectonism in the Ecuadorian northwestern cordil‐
lera: Metallogenic and Geodynamic implications. PhD thesis, Katholieke Universiteit Leuven, 150 p. 38 Wendt, I, & Carl, C. (1991): The statistical distribution of the mean squared weighted deviation. Chemical geology 86; 275‐285. Winkler, W., Villagomez, D., Spikings, R., Abegglen, P, Tobler, S, Eguez, A. (2005): The Chota basin and its significance for the inception and tectonic setting if the inter‐Andean depression in Ecuador. Journal of South American Earth Sciences 19; 5‐19. Zellmer, G. F. (2008): Some first‐order observations on magma transfer from mantle wedge to upper crust at volcanic arcs. In: Annen, C. & Zellmer, G. F. (eds), Dy‐
namics of Crustal Magma Transfer, Storage and Differ‐
entiation. Geological Society, London, Special Publica‐
tions 304; 15‐31. Appendix I – Accommodation of convergence obliquity at the Ec‐
uadorian margin throughout the Tertiary The occurrence of non‐orthogonal plate conver‐
gence implies two end‐member geodynamic situations where convergence obliquity is either accommodated by oblique subduction slip or overriding plate strain partitioning possibly in‐
cluding forearc sliver displacement (McCaffrey 1992). Whereas the former is likely to influence asthenospheric flow dynamics, the latter consti‐
tutes an important factor for crustal magma as‐
cent and pluton emplacement mechanisms (e.g., Glazner 1991). Therefore, the mode of conver‐
gence obliquity accommodation at a convergent plate margin has to be critically evaluated in or‐
der to discuss potential controls on overriding plate magmatism (cf. Appendix II). Ego et al. (1996) show that the present‐day Ecua‐
dorian margin is weakly decoupled where 75‐90% of the convergence obliquity are transferred into trench‐oblique, due east (87°) subduction slip, whereas only 10‐25% obliquity are accommo‐
dated by the upper plate. Current dextral forearc sliver displacement rates along the CPPF can ac‐
count for most of the northwards increasing up‐
per plate trench‐parallel component (Ego et al. 1996), although a number of additional Andean‐
trending fault systems (including the CTSZ and the Peltetec fault in Ecuador) also accommodate convergence obliquity and currently act as a re‐
gional restraining bend (Winkler et al. 2005). The highly oblique subduction slip of the Ecuadorian‐
Colombian margin is uncharacteristic for the southern‐central Andes where subduction is mostly trench‐normal; Sébrier & Bellier (1993) suggest that the allochthonous, rheologically strong oceanic plateau basement of the northern Andes might control the degree of margin de‐
coupling of this regional arc segment. Using slip rates comparable to Pliocene‐present‐
day values Hungerbühler et al. (2002) obtain a total displacement estimate of 100‐130 km along the CPPF since the Mid‐Miocene which they use to perform a palinspastic forearc sliver recon‐
struction. These authors show that in the recon‐
structed setting the Manabi and Progreso forearc basins are juxtaposed against Interandean re‐
gions of Miocene marine transgressions; these depositional environments are correlatable in terms of a continuous proximal‐distal basin fa‐
cies, thus indicating that the scale of their dis‐
placement estimate is accurate. Consequently, using CPPF slip rates as a proxy for the oblique convergence component accommodated by the upper plate, one would assume that the Ecuador‐
ian margin must have been constantly weakly decoupled since the Mid‐Miocene. However, the direction of plate convergence at the Ecuadorian margin has been almost due east since 28.3 Ma (Chapter 2: Fig. 6), paralleling the present‐day subduction slip (instead of the pre‐
sent‐day direction of convergence). If present‐day subduction slip directions were applied, the cen‐
tral Ecuadorian margin would be almost com‐
pletely coupled during most of the Late Oligo‐
cene‐Miocene, and weakly to moderately de‐
coupled only from 20.2‐16.0 Ma and from 4.9 Ma until the present day. Significant convergence obliquity accommodated by major forearc sliver displacement would thus be restricted to these periods. Clearly, these considerations are in dis‐
agreement with the expected upper plate strike‐
slip displacement from the reconstruction of Hungerbühler et al. (2002), as displacing the forearc sliver for 100‐130 km since only 4.9 Ma required unrealistically high displacement rates, which would be difficult to reconcile with Pleisto‐
cene values (Trenkamp et al. 2002). Consistent evidence for strain partitioning with a significant dextral strike‐slip component along the repeatedly reactivated major fault systems below the Western Cordillera during the Tertiary exists, although the exact time periods are not always well constrained. Dextral shearing along the CTSZ in northern Ecuador has a minimum age of 48.3±0.6 Ma (hornblende K‐Ar age of foliated diorite emplaced within the CTSZ; Hughes & Pi‐
latasig 2002). A short distance to the west of the CTSZ, Chiaradia et al. (2008) show S‐C fabrics in altered rocks associated with the Macuchi‐hosted La Plata ore deposit; these authors interpret the structures as indicative of Late Eocene dextral 39
transpression related to distributed shearing along the CTSZ. Mid‐Late Eocene whole‐scale dextral strike‐slip motion of the Ecuadorian forearc sliver along the proto‐CPPF and CTSZ fur‐
ther produced intra‐forearc block rotations where individual blocks are separated by a set of NNW‐trending sinistral dip‐slip faults (Daly 1989). The whole structure of the Western Cordillera is interpreted as a positive flower structure formed in response to dextral transpressional reactiva‐
tion of the 35°E‐dipping suture zone forming the root of the CPPF, as well as the deeply‐rooted CTSZ (Guillier et al. 2001, and references therein). Finally, Winkler et al. (2005) infer that dextral transpression along a restraining bend (including the CTSZ, CPPF, Peltetec fault, and Chingual‐La Sofia fault in Ecuador) has been active since at least 15 Ma, with a pulse of increased activity since 6 Ma being responsible for the inception of the IAD as a full‐, and locally half‐ramp basin. Four possible explanations for the discrepancy between convergence obliquity‐based predicted limited strike‐slip motion since the Mid‐Miocene and upper plate geological evidence for more significant strike‐slip motion exist, and might all be applicable to some degree: (1) Computation of Somoza’s (1998) rotation poles for the Ecuadorian margin (Chapter 2: Fig. 6) slightly underestimates convergence obliquity. However, using convergence parameters calcu‐
lated from Pardo‐Casas & Molnar (1987) instead yields an even less oblique direction of conver‐
gence with respect to the present‐day subduction slip. While Somoza (1998) does not present any rotation pole uncertainties, the uncertainty ranges shown by Pardo‐Casas & Molnar (1987) indicate that the precision of the rotation poles is insufficient to further evaluate temporal varia‐
tions in obliquity at this time scale, as was al‐
ready noted by Daly (1989). Generally, recon‐
structed convergence parameters always average plate motion over multi‐m.y. periods such that any short‐term variations are smoothened out and thus difficult to detect. (2) The Tertiary paleo‐subduction slip deviated from the direction of the present‐day subduction slip. Considering a multi‐m.y. lag time, this would be a likely response to changes of the conver‐
gence direction, especially given the proximity of 40 the Ecuadorian‐Colombian margin to the Nazca‐
Cocos‐Caribbean plate boundary to the north, the transition to the northern Peruvian flat slab segment to the south, and the potential influence of Carnegie Ridge subduction, which all likely in‐
creases the Miocene subducting slab geometric complexity (e.g., Gutscher et al. 1999; Taboada et al. 2000). The complex margin geometry and out‐
crop pattern of Tertiary arc units in Ecuador (Chapter 2: Figs. 1; 2) prevents an arc geometry‐
based discussion of paleo‐subduction slip direc‐
tions even if, for simplicity, subvertical translitho‐
spheric magma ascent was assumed. Overall, one expects a regionally more homogenous subduc‐
tion geometry along the NW South American margin prior to the Late Oligocene Farallon plate fission. (3) The Ecuadorian margin and trench geometry changed through time; the margin geometry could obviously be strongly influenced by forearc sliver displacement. Daly (1989) estimates a minimum clockwise trench rotation of 20° since the Oligocene, but does not present any argu‐
ments in favor of this estimate. (4) The time‐dependent correlation of sedimen‐
tary basin facies by Hungerbühler et al. (2002) based on continuous displacement since the Mid‐
Miocene is inaccurate and overestimates the to‐
tal forearc sliver displacement in the latest Terti‐
ary; instead, forearc sliver displacement of 100‐
130 km along the proto‐CPPF represents cumula‐
tive, non‐continuous displacement processes throughout the whole Tertiary (Witt et al. 2006). In conclusion, if the present‐day mode of conver‐
gence partitioning at the Ecuadorian margin holds some significance for the Tertiary subduc‐
tion system, and at the same time substantial forearc sliver displacement took place since the Mid‐Miocene, it seems likely that the margin ge‐
ometry and the degree of margin decoupling, i.e., the component of upper plate accommodation of oblique plate convergence, varied through time. Therefore, overriding versus downgoing plate strain partitioning in response to changing con‐
vergence obliquities cannot be reliably predicted with currently available data, but geological evi‐
dence for multiple phases of Tertiary strike‐slip deformation in the overriding plate exists. From the Eocene to the Late Oligocene, plate conver‐
gence was significantly more oblique with respect to the present‐day subduction slip (Chapter 2: Fig. 6) making short‐term variations in subduction slip less likely, such that consistently increased margin decoupling was a likely consequence. These considerations are in general agreement with the notion of Witt et al. (2006) discussed above, and would thus be equally applicable if the Mid‐Miocene sedimentary basin facies corre‐
lation of Hungerbühler et al. (2002) was incor‐
rect. References Chiaradia, M., Tripodi, D., Fontboté, L., Reza, B. (2008): Geologic setting, mineralogy, and geochemistry of the Early Tertiary Au‐rich volcanic‐hosted massive sulfide deposit of La Plate, Western Cordillera, Ecuador. Eco‐
nomic Geology 103; 161‐183. Daly, M. C. (1989): Correlation between Nazca/Farallon plate kinematics and forearc basin evo‐
lution in Ecuador. Tectonics 8:769–790. Ego, F., Sébrier, M., Lavenu, A., Yepes, H., Egues, A. (1996): Quaternary state of stress in the Northern An‐
des and the restraining bend model for the Ecuadorian Andes. Tectonophysics 259; 101‐116. Glazner, A. F. (1991): Plutonism, oblique subduction, and continental growth: an example from the Meso‐
zoic of California. Geology 19; 784‐786. convergent subduction: the Andean case. Extended conference abstracts, ISAG 1993, Oxford, 139‐142. Somoza, R. (1998): Updated Nazca (Farallon)—South America relative motions during the last 40 My: impli‐
cations for mountain building in the central Andean region. J S Am Earth Sc 11; 211‐215. Taboada, A., L. A. Rivera, A. Fuenzalida, A. Cisternas, H. Philip, H. Bijwaard, J. Olaya, and C. Rivera (2000), Geo‐
dynamics of the northern Andes: Subductions and intracontinental deformation (Colombia), Tectonics, 19(5), 787–813. Trenkamp, R., Kellog, J. N. Freymueller, J. T. Mora, H. P. (2002): Wide plate margin deformation, southern Central America and northwestern South America, CASA GPS observations: Journal of South American Earth Sciences, v. 15, p. 157‐171. Winkler, W., Villagomez, D., Spikings, R., Abegglen, P, Tobler, S, Eguez, A. (2005): The Chota basin and its significance for the inception and tectonic setting if the inter‐Andean depression in Ecuador. Journal of South American Earth Sciences 19; 5‐19. Witt, C., J. Bourgois, F. Michaud, M. Ordoñez, N. Jimé‐
nez, and M. Sosson (2006): Development of the Gulf of Guayaquil (Ecuador) during the Quaternary as an ef‐
fect of the North Andean block tectonic escape, Tec‐
tonics, 25, TC3017, doi:10.1029/2004TC001723. Guillier, B. Chatelain J.L. Jaillard, E., Yepes, H., Poupinet, G., Fels, J.F. (2001): Seismological evidence on the geometry of the orogenic system in central‐
northern Ecuador (South America): Geophysical Re‐
search Letters, v. 28, p. 3749‐3752. Hughes R. A., Pilatasig L. F. (2002): Cretaceous and Tertiary terrane accretion in the Cordillera Occidental of the Andes of Ecuador. Tectonophysics 345:29–48. Hungerbühler, D., Steinmann, M., Winkler, W., Sew‐
ard, D., Egüez, A., Peterson, D. E., Helg, U., Hammer, C. (2002): Neogene stratigraphy and Andean geodynam‐
ics of southern Ecuador. Earth Science Reviews 57; 75–
124. McCaffrey, R. (1992): Oblique Plate Convergence, Slip Vectors, and Forearc Deformation, J. Geophys. Res., 97(B6), 8905–8915. Pardo‐Casas, F. & Molnar, P. (1987): Relative Motion of The Nazca (Farallón) and South American Plates since Late Cretaceous Time. Tectonics 6; 233‐248. Sébrier, M. & Bellier, O. (1993): How is accommodated the parallel‐to‐the‐trench slip compo‐nent in oblique 41
Appendix II – Overriding plate structural controls on the spatio‐
temporal distribution of Tertiary plutons in Ecuador Overriding plate tectonics, stress regime and arc magmatism show coupled behavior on a regional to local scale: strike‐slip deformation as a result of strain partitioning in the overriding plate is mainly localized in the rheologically weak arc magmatic zone (Dewey et al. 1998). Melt ascent and emplacement in the crust can be efficiently focused by structures and is significantly aided by deformation (Saint Blanquat et al. 1998; Vigner‐
esse & Clemens 2000), and may in turn induce further strike‐slip partitioning (Saint Blanquat et al. 1998). Detailed recent reviews of the complex mechanisms of arc magma ascent and emplace‐
ment are presented by Richards (2003) and Cem‐
brano & Lara (2009). In the following, we focus on investigating whether and how the spatial and temporal distribution of Tertiary plutonism in Ecuador is affected by: (1) the regional signifi‐
cance of structural control for localizing intru‐
sions; and (2) the role of synintrusive deforma‐
tion and the local‐regional stress field on pluton emplacement. Plutons mapped as Tertiary intru‐
sions in the Eastern Cordillera have Late Creta‐
ceous‐Early Tertiary K‐Ar ages which are poten‐
tially thermally disturbed on a regional‐scale (Peltetec event; Litherland et al. 1994), and are thus excluded from any further discussion here. A few general considerations apply for these two points. Due to intrinsic overpressuring arc magma is principally able to ascend and reach the Earth’s surface irrespective of the prevailing local‐
regional stress regime (e.g., Paterson & Fowler 1993; Saint Blanquat et al. 1998; Cembrano & Lara 2009); a compressional stress component does not prevent magma from ascending along pre‐existing or newly formed structures but may in fact enhance magma ascent by tectonic over‐
pressuring (especially at lower crustal levels; Saint Blanquat et al. 1998). Local transpressional or transtensional stress settings (e.g., associated with restraining or releasing bend geometries of large strike‐slip systems; Sylvester 1988) prefer‐
entially localize rapid magma ascent, because 42 they commonly provide subvertically oriented high‐permeability structures (e.g., Richards 2003). However, particularly in the upper crust, ascending magma may use any available struc‐
tural weakness such that ascent is not necessarily vertical (Saint Blanquat et al. 1998; Kalakay et al. 2001). Furthermore, the stress regime prevailing at a given time does not control the orientation of pre‐existing structures (Cembrano & Lara 2009). Space creation for intrusion emplacement in the upper crust involves displacement of the crustal host rocks, either by means of deforma‐
tion where fault kinetics control emplacement rates (Glazner 1991; Grocott et al. 1994; Acocella et al. 2008) or/and by ballooning and roof uplift (Paterson & Fowler 1993; Saint Blanquat et al. 2006). Thus, emplacement of intrusions along pre‐existing or newly formed structures com‐
bined with synintrusive deformation does not seem to be a general requirement for crustal magma ascent or pluton emplacement, but these processes are expected to be positively corre‐
lated, and possibly feedback‐related with each other. The spatial distribution of Tertiary intrusions in Ecuador closely follows the major NNE‐ and ESE‐
trends of upper plate structures (Fig. 2). In this context, plutons are not expected to be directly localized along major first‐order structures but should rather be emplaced in associated periph‐
eral areas of dilation (Richards 2003). Major in‐
trusive belts center on deeply‐rooted faults and translithospheric suture zones between the mainland and the allochthonous western oceanic domain or the southern Amotape terrane, re‐
spectively: 
The N‐ to NE‐trending CTSZ at the western flank of the Western Cordillera is spatially associated and aligned with the northern batholiths of Santiago and Apuela‐Nanegal, the central Ecuadorian batholith system (Corazon, Telimbela‐Chazo Juan, Balsa‐
pamba‐Las Guardias, Echeandia‐Industria) as well as multiple small intrusions. Until recently, the CTSZ used to be regarded as a suture zone (Hughes & Pilatasig 2002) al‐
though Vallejo et al. (2006, 2009) proposed a geodynamic model for the Ecuadorian margin where the Macuchi island arc is autochthonous dismissing the suture origin of the CTSZ. Nonetheless, the great strike length of the shear zone and seismic stud‐
ies (Guillier et al. 2001) indicate that the structure extends to deep crustal (possibly transcrustal) levels, although it might not be a translithospheric transform fault such as the CPPF. As described above (Appendix II of Chapter 2), geologic evidence suggests multiple phases of dextral transpressional movement along the shear zone since the Eocene. 
The NE‐ to NNE‐trending CPPF represents the western limit of a Late Cretaceous su‐
ture zone below the Western Cordillera (e.g., Vallejo et al. 2006); it has been reacti‐
vated as the present‐day CPPF with a dex‐
tral transpressional sense of movement and accommodates a significant compo‐
nent of the present‐day convergence obliquity (Ego et al. 1996). It is spatially as‐
sociated with the Chaucha batholith and multiple smaller intrusions at its southern end where it splays off and intersects the Western Cordillera towards the Gulf of Guayaquil. The preferential emplacement of Tertiary intrusions along the CPPF and CTSZ was already noted by Litherland & Aspden (1992), although the tectonomag‐
matic model presented by these authors does not withstand modern concepts of arc magma genesis (e.g., Stern 2002; Richards 2003). 
The Cangrejos‐Zaruma intrusive belt occu‐
pies a central axial position between the Piñas‐Portovelo and Jubones fault systems, probably in the vicinity of the northern limit of Amotape basement (Litherland et al. 1994). Towards the west, the Piñas‐
Portovelo fault joins with the La Palma‐El Guayabo and Tahuin Dam (Naranjos) faults; these faults delimit the deeply exhumed metamorphic Raspas complex whose struc‐
tural position has been related to the an‐
cient Amotape suture zone (Bosch et al. 2002). Thus, while the exact location of the northern Amotape suture is concealed be‐
low Tertiary cover sequences, the Cangre‐
jos‐Zaruma belt seems to be emplaced in a proximal and subparallel position with re‐
spect to the suture zone. Further south, the voluminous Portachuela batholith is em‐
placed and exhumed along the N‐S trend‐
ing Las Aradas fault between the Amotape terrane and the Eastern Cordillera (Lither‐
land et al. 1994). The Andean‐trending CTSZ and CPPF show a sub‐
parallel orientation with respect to the central‐
northern Ecuadorian margin and the downgoing slab, and are thus potentially favorably aligned with zones of asthenospheric partial melting and lithospheric magma ascent. In contrast, the Can‐
grejos‐Zaruma intrusive belt forms a transverse structure with respect to the margin trend, char‐
acterized by roughly coeval (on a time scale of a few m.y.) pluton emplacement at its eastern and western end and, by inference, along its whole strike length. While the present‐day slab geome‐
try below this arc segment is likely contorted or discontinuous as a result of the flat‐steep slab transition between northern Peru and central Ecuador (Gutscher et al. 1999), such geometric complexities are unlikely to have existed in the Late Oligocene‐Early Miocene when most of the intrusions formed; a continuous subduction of the Farallon plate can be assumed for that time. Slab geometry and the orientation of the zone of asthenospheric partial melting alone thus fail to explain the intrusive belt alignment, such that the arc‐transverse pluton emplacement trend must have been significantly structurally controlled. The concentrated occurrence of Oligocene‐
Miocene plutons along the flanks of the CTSZ suggests this structure in part controlled magma ascent and pluton emplacement. Subvertical transcrustal structures are expected to channel‐
ize ascending magma (e.g., Cembrano & Lara 2009). However, at c. 1°30’ S latitude, Guillier et al. (2001) show present‐day seismicity patterns defining 35°E dipping planes at mid‐ to deep crustal levels which these authors interpret as the traces of the CTSZ and CPPF fault planes, as they intersect these structures at surface levels. Reactivation of the faults caused deformation of their upper portions resulting in subvertical dips close to the surface (Guillier et al. 2001). If a simi‐
lar, non‐vertical dip had already been established during the Oligocene‐Miocene this scenario could be interpreted in the following ways: (1) an addi‐
tional deeper subvertical portion of the CTSZ ex‐
ists and was exploited by ascending magmas, but 43
is not seismically active at present; (2) the proto‐
CTSZ only controlled pluton emplacement (by tectonic space creation) at shallow crustal levels, but did not significantly influence magma ascent at depth; (3) significant non‐vertical magma as‐
cent along the proto‐CTSZ at deep‐ to mid‐crustal levels took place. Given the deep crustal nature of the CTSZ, we expect it to be principally able to efficiently channelize magmas through the whole crust, in particular as the mafic‐ultramafic oce‐
anic basement units of the Western Cordillera are rheologically strong and require a high differen‐
tial stress to fracture. Repeated transpressional reactivation of the shear zone throughout the Tertiary might have aided magma ascent by tec‐
tonic overpressuring. Non‐vertical magma ascent along thrust ramps has been documented in a number of settings and can be a viable mecha‐
nism for magma ascent in an overall transpres‐
sional‐compressional stress regime (e.g., Kalakay et al. 2001). If magma ascent in the northern Ec‐
uadorian arc segment during the Oligocene‐
Miocene was partly controlled by east‐dipping thrust geometries, the zone of partial melting at depth could extend well to the east of the pre‐
sent‐day CTSZ, in agreement with the broad landwards extent of arc magmatism in the south‐
ern‐central Ecuadorian arc segments further south. While detailed kinematic structural studies of Tertiary pluton emplacement in Ecuador’s West‐
ern Cordillera are lacking, it can be inferred from the discussion above that distributed shear re‐
lated to forearc sliver displacement along the CPPF and CTSZ, combined with strike‐slip reacti‐
vation of the older suture zones further east (Litherland et al. 1994; Winkler et al. 2005) has been intermittently active throughout the Terti‐
ary, and probably resulted in partially syntectonic intrusive activity. Descriptive studies available for the major intrusions of the Western Cordillera are summarized in Prodeminca (2000a) and gen‐
erally support this notion. Prodeminca (2000a) note that most intrusions are spatially associated with second‐order NE‐ to ENE‐trending faults which under dextral transpression should pro‐
duce local dilation. In the northern Western Cordillera voluminous magmatism of the Apuela‐Nanegal batholith might be related to its structural position be‐
44 tween two major structures, the CPPF and the CTSZ, and second‐order lineaments associated with these structures (but see discussion on addi‐
tional across‐arc lineaments in Chapter 2). While the batholithic intrusions in central Ecuador (Balsapamba‐Las Guardias, Telimbela‐Chazo Juan, plus associated intrusions) are spatially associ‐
ated with lineaments of various orientations mostly trending subparallel to the CTSZ (Prodeminca 2000a), none of these have been mapped as faults on regional Western Cordillera maps (Fig. 1; McCourt et al. 1998; Hughes et al. 1998). At the southern end of the Western Cordil‐
lera, Chaucha batholith emplacement is inferred to be generally related to the Bulubulu fault which forms part of the CPPF; a number of asso‐
ciated NE‐ and NW‐trending faults are thought to have controlled individual intrusion emplacement and porphyry mineralization (Prodeminca 2000a). Rapid unroofing of the relatively young Chaucha batholith was associated with regional contrac‐
tion leading to basin inversion in the Interandean region at c. 9 Ma (Hungerbühler et al. 2002). In the Zaruma region close to the eastern end of the Late Oligocene‐Early Miocene Cangrejos‐
Zaruma intrusive belt, growth sequences of intru‐
sion‐hosting Saraguro Group volcanics form thickening wedges towards the southern Piñas‐
Portovelo fault, indicative of synvolcanic normal fault slip (Spencer et al. 2002). Further north, a normal slip component, albeit of unconstrained age, is detected at the E‐W‐trending Jubones fault (Litherland et al. 1994). These observations are in agreement with Steinmann’s (1997) pro‐
posal of regional horizontal extension during the Oligocene‐Early Miocene deposition of Saraguro Group ignimbrites, which are inferred to have been sourced from fissure eruptions and caldera‐
forming events. Horizontal extension in southern Ecuador was followed by transpression which is recorded by inversion of the Piñas‐Portovelo fault and folding in the area north of the fault producing a major anticline subparallel to the Cangrejos‐Zaruma intrusive belt (Spencer et al. 2002), as well as by a conjugate set of NE‐trending faults with evidence for dextral movement (Prodeminca 2000a). Fur‐
thermore, whole‐scale tilting of the Saraguro Group volcanic sequence north of the Piñas‐
Portovelo fault is observed (now dipping 30° to the SW; Spencer et al. 2002). Plutons north of the Piñas‐Portovelo fault which, based on the radio‐
metric ages obtained in this study, can be in‐
ferred to be of mainly Early Miocene age, show asymmetric sigmoidal plan‐view geometries in‐
dicative of syntectonic intrusion into a dextral transpressional stress field. Further, Au‐bearing hydrothermal quartz‐calcite veins in the Zaruma‐
Portovelo mining district are related to NW‐
striking faults moderately dipping to the SW; these faults show S‐C fabrics and shear banding, in agreement with vein formation under dextral transpression. The veins are interpreted as partly originating from magmatic fluids thus necessitat‐
ing roughly coeval and hence syntectonic mag‐
matism (Spencer et al. 2002). These considera‐
tions indicate that Late Oligocene‐Early Miocene magmatism forming the Cangrejos‐Zaruma intru‐
sive belt is at least in part syntectonic in nature. Transpressional deformation in this region might either be related to the oblique subduction set‐
ting, or, possibly, to the post‐Paleocene 25±12° clockwise block rotation inferred for the Amo‐
tape terrane from paleomagnetic studies (Mi‐
touard et al. 1990). There does not seem to be a first‐order relation‐
ship between the intensity of plutonism, in par‐
ticular the significant increase in the Late Oligo‐
cene, and changes of the regional stress regime in Ecuador. Steinmann (1997) and Hungerbühler et al. (2002) infer a regional tensional stress field from 40‐20 Ma, with a compressional pulse at c. 19 Ma, followed by another period of horizontal tensional stress in the Interandean region from 15‐11 Ma, and compression from 9‐8 Ma. How‐
ever, compared to their detailed studies of Mid‐
Late Miocene sedimentary basins, the geologic evidence presented by these authors to constrain the Oligocene‐Early Miocene stress field is rather limited, as it is solely based on extensional forearc deformation at that time described by Daly (1989), combined with an inferred cause‐
effect relationship of regional extension and Saraguro Group ignimbrite eruption. Conse‐
quently, the lack of correlation between varia‐
tions in the regional stress field and the peak in shallow crustal arc magmatism initiating in the Late Oligocene‐Early Miocene might be due to insufficient knowledge of the paleo‐stress field at the Ecuadorian margin. Overall, however, it is more likely that plutonic activity increased con‐
comitant with volcanism during the Late Oligo‐
cene to Mid‐Miocene as part of the flare‐up event in arc magmatism discussed in Chapter 2, for which larger geodynamic controls are ulti‐
mately inferred. On a local scale, Prodeminca (2000a) note that periods of intense magmatism and mineralization tend to be associated with inferred changes in the local stress regime, particularly at the onset of post‐compressional tensional periods. As dis‐
cussed above, however, local variations in stress regime might in part reflect coupling between magmatic and tectonic processes such that local extension might in part be induced by intrusive activity. Clearly, detailed structural studies of Ter‐
tiary intrusions hosted by the Western Cordillera are needed to further discuss the relationships between plutonism and the regional and local tectonic environment. Significant Tertiary forearc sliver displacement implies that the plutons of the Western Cordillera intruded at more southern latitudes than their present‐day position, relative to the Ecuadorian mainland and the “fixed” Tertiary intrusions hosted by the southern‐central Ecuadorian arc segments. If the forearc sliver displacement esti‐
mate of 100‐130 km since the Mid‐Miocene by Hungerbühler et al. (2002) is correct, most intru‐
sions will have undergone significant whole‐scale latitudinal displacement, as their timing of em‐
placement predates the displacement period. Alternatively, following the reasoning of Witt et al. (2006), latitudinal displacement could increase with pluton age significantly beyond the Mid‐
Miocene. A forearc sliver displacement of 100‐
130 km since the Mid‐Miocene implies juxtaposi‐
tion of the Oligocene‐Early Miocene batholithic intrusions of central Ecuador with the volumi‐
nous, slightly younger intrusions in the Chaucha area at the limit between the northern and cen‐
tral Ecuadorian arc segments. The concentration of intrusive activity in this presently highly tec‐
tonized region could reflect a concentration of strike‐slip deformation potentially favorable for crustal magma ascent, space creation, and thus pluton emplacement (e.g., Glazner 1991; Rich‐
ards 2003). Alternatively, or in addition to the preceding point, a positive asthenospheric heat anomaly might have locally persisted below the 45
region of the central Ecuadorian arc segment, further accentuating (or causing?; cf. discussion in Chapter 2) the general peak in arc magma pro‐
ductivity in the Late Oligocene to Mid‐Miocene. In conclusion, Tertiary intrusions are preferen‐
tially localized along deeply‐rooted fault zones, most of them sutures, between the Ecuadorian mainland and the Amotape terrane, or the alloch‐
thonous forearc sliver floored by oceanic plateau basement, suggesting a fundamental control of these structures on magma ascent and pluton emplacement. Overriding plate deformation in response to oblique plate convergence and par‐
tial decoupling of the Ecuadorian margin was mainly transpressional‐transtensional, and is ex‐
pected to have been intermittently active throughout large parts of the Tertiary until the present day. Transpressional reactivation of ma‐
jor fault zones and the associated distributed shear might have aided magma ascent and con‐
tributed to space creation for intrusion em‐
placement. The inferred mid‐deep crustal thrust geometry of the CTSZ and CPPF (Guillier et al. 2001) implies that magma ascent along these structures was not necessarily subvertical. The high spatial concentration of cogenetic Oligo‐
cene‐Miocene intrusions and volcanics in the central Ecuadorian arc segment, bracketed be‐
tween two suture zones, might be related to con‐
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47
Appendix III – Accuracy of pub‐
lished K‐Ar (and ZFT) ages of Ter‐
tiary intrusions in Ecuador Before the present work the geochronologic framework on the timing of Tertiary plutonism in Ecuador exclusively relied on K‐Ar (mostly biotite and hornblende fractions, or whole‐rock samples) plus few ZFT datations. An exception is Bineli Betsi (2007) who presents two zircon U‐Pb ages of intrusions spatially associated with the Rio Blanco low sulfidation epithermal deposit in the central Western Cordillera of Ecuador. The clo‐
sure temperature range (e.g., Chesley 1999) for the K‐Ar isotopic systems of hornblende (490‐
570°C) and, particularly, biotite (260‐350°C) makes these minerals susceptible to thermal dis‐
turbance by either burial, proximal emplacement of younger intrusions, or, especially, by porphyry‐
related hydrothermal alteration where initial fluid temperatures of >500°C for early potassic alteration are common (e.g., Seedorff et al. 2005). This issue is even more pertinent for whole‐rock K‐Ar ages, where a reliable closure temperature is difficult to estimate, or for ZFT data (closure temperature 260±25°C; Foster et al. 1996). A field example illustrating this effect is the small El Tingo pluton in southern Ecuador with K‐Ar hornblende and biotite ages of 47‐50 Ma (Kennerley 1980) where Hungerbühler et al. (2002) obtained a Miocene ZFT age of 21±3 Ma. Zircon U‐Pb age data acquired on Late Tertiary intrusions allow us to evaluate whether previ‐
ously reported ages obtained by these methods represent magmatic cooling, or whether they have been thermally disturbed, thus partly offset towards a younger intrusive or hydrothermal event, or pluton exhumation. Furthermore, most K‐bearing minerals are highly susceptible to al‐
teration, and already small amounts of secondary replacement minerals may suffice to produce K‐
Ar ages of uncertain geological significance, an issue which also pertains to Ar‐Ar ages (Snee 2002). As different types of hydrothermal altera‐
tion of variable intensity are a common feature of most intrusive complexes of Ecuador's Western Cordillera (e.g., Prodeminca 2000a), the accuracy of K‐Ar ages can be tested with a relatively altera‐
tion‐resistant mineral such as zircon. 48 Hornblende K‐Ar ages of granodiorites of the Apuela‐Nanegal batholith are 18.5±0.9 Ma (Van Thournout 1991), 16.5±0.8 Ma (Prodeminca 2000a), and 14.5±0.2 Ma (MMAJ/JICA 1992), with additional biotite K‐Ar ages of 16.0±0.8 Ma (Prodeminca 2000a), 15.8±0.6 Ma (Van Thournout 1991) and 13.0±0.6 Ma (MMAJ/JICA 1992). The latter age, reported by MMAJ/JICA (1992) for the Cuellaje prospect area, is identical within error with our zircon U‐Pb age of 12.87±0.08 Ma on the same lithology, i.e., por‐
phyry‐hosting hornblende‐biotite‐bearing grano‐
diorite, and is thus inferred to represent a mag‐
matic cooling age. This indicates that the young‐
est pulse of batholith emplacement at Cuellaje occurred at relatively shallow crustal levels at wall rock temperatures <260‐350°C (the biotite K‐
Ar closure temperature range). Batholith mag‐
matic cooling rates were very high as evidenced by identical zircon and biotite ages; subsequent burial and reheating to temperatures >260‐350°C did not take place. Similarly, we consider the older hornblende K‐Ar ages referred to above to define distinct pulses of batholith emplacement at shallow crustal levels, since the K‐Ar horn‐
blende and biotite ages of MMAJ/JICA (1992) are identical within error. Geochronologic data thus suggest a composite, multi‐intrusive nature of the batholith which is in agreement with detailed geologic and petrographic investigations by Sala‐
zar (2007), defining various, often fault‐bounded batholith lithologies including tonalite, quartz monzonite, quartz‐diorite and monzogranite, in addition to a major granodiorite lithology. Age data for the multiple porphyry stocks and dikes intruding the batholith define age ranges of 11.1‐8.8 Ma for the porphyries at Cuellaje (whole rock K‐Ar; MMAJ/JICA 1992) and 7.9‐5.6 Ma at Junin (whole rock K‐Ar, MMAJ/JICA 1992; biotite‐
hornblende K‐Ar, Prodeminca 2000a). While we cannot assess the accuracy of the Cuellaje ages, ages reported for the Junin porphyries, classified as "quartziferous porphyry" by MMAJ/JICA (1992) and Prodeminca (2000a), are significantly younger than our hornblende granodiorite por‐
phyry dike zircon U‐Pb age of 9.01±0.05 Ma. The lithological classification as "quartziferous por‐
phyry" used by MMAJ/JICA (1992) and Prodeminca (2000a) is somewhat ambiguous as evidenced by their reported 2 m.y. K‐Ar age dif‐
ference for "quartziferous porphyry" exceeding analytical errors, and we infer that these K‐Ar ages date cooling or (partial) thermal resetting of one or more younger porphyry intrusive events and/or their associated hydrothermal systems. Field evidence indicates at least three major and several minor porphyry phases are present at the Junin prospect (Salazar 2007). The 9 Ma horn‐
blende granodiorite porphyry age presented in this study seems to define the earliest timing of porphyry emplacement, but older pulses of activ‐
ity are likely to have occurred, as indicated by a 9.5±0.2 Ma zircon antecryst. Younger pulses ap‐
parently center on the 7.9‐7.3 Ma and 6.1‐5.6 Ma time ranges; further resolving the timing of these multiple intrusive events is complicated by petro‐
graphic evidence of high‐T potassic alteration overprinted by pervasive phyllic alteration (Sala‐
zar 2007). Thermal disturbance is evidenced by a 7.5±0.2 Ma biotite K‐Ar age of altered host granodiorite (Prodeminca 2000a). Continuous multi‐m.y. magmatism focused on the Junin area might generally have led to a local shallow crustal heat anomaly, such that the young K‐Ar ages may not be used as accurate estimates for intrusive events of the Junin porphyry stocks and dikes on a local scale, although they may be broadly used as proxies for the latter. A further discussion of the Junin porphyry system is provided in Chapter 3, where two Re‐Os molybdenite ages (6.13±0.03 Ma and 6.63±0.04 Ma) are presented, which sup‐
port the occurrence of several younger porphyry intrusive events as inferred from K‐Ar data. Previous studies based on hornblende and biotite K‐Ar ages, and U‐Pb data presented in this work define two age clusters for the Balsapamba‐Las Guardias batholith in central Ecuador. In the Las Guardias area an earlier peak comprises three hornblende and biotite ages ranging from 34.3 to 30.1 Ma (MMAJ/JICA 1989; McCourt et al. 1997). In addition, Henderson (1979) reports ages of 30.8±1 Ma (K‐Ar biotite) and 19.2±3 Ma (K‐Ar hornblende) on a quartz‐diorite sample from the same area where he prefers the latter as esti‐
mate for the age of pluton emplacement. These ages were recalculated by Kennerley (1980) using decay constants of Steiger & Jäger (1977) to 31.7±1 Ma and 19.8±3 Ma. The fact that his hornblende K‐Ar age is significantly younger than the biotite K‐Ar age of the same sample suggests hydrothermal alteration affected the biotite and/or hornblende thus rendering these ages potentially inaccurate. In the Balsapamba area, K‐
Ar hornblende ages of 33.1±0.4 and 25.7±0.9 Ma are reported (MMAJ/JICA 1989; McCourt et al. 1997). There, the younger age cluster is more dominant including our new zircon U‐Pb ages at 21.46±0.08 Ma (batholith) and 21.22±0.17 Ma (porphyry dike), and hydrothermal molybdenite Re‐Os ages of 19.9±0.3 Ma (Chiaradia et al. 2004) and 21.5±0.1 Ma (Chapter 3) related to the Balsapamba porphyry system. Preservation of biotite ages of the older age clus‐
ter indicates the batholith has resided at shallow crustal levels below the biotite K‐Ar closure tem‐
perature since then. Furthermore, as a systematic younging from hornblende to biotite K‐Ar ages within the older age cluster is not observed, and biotite K‐Ar ages show intra‐cluster scattering outside of their reported uncertainty range, initial batholith emplacement should have taken place at relatively shallow crustal levels where the age scatter reflects magmatic cooling from multiple intrusions and/or hydrothermal resetting. A shal‐
low crustal emplacement environment at least for the younger part of the batholith is in agree‐
ment with our field observations of intrusive brecciation along 21.2 Ma porphyry dike margins, requiring brittle deformation mechanisms. The 19.9±0.3 Ma molybdenite Re‐Os age reported by Chiaradia et al. (2004) is indicative of the addi‐
tional presence of a younger, post‐21.2 Ma por‐
phyry intrusion at Balsapamba, since the maxi‐
mum estimate for the lifetime of a large, single intrusion‐driven hydrothermal system is 0.8 m.y. (Cathles et al. 1997), and typical lifetimes are sig‐
nificantly shorter still (e.g., Shinohara & Heden‐
quist 1997). This is in agreement with published lithological data indicating a composite nature of the batholith (MMAJ/JICA 1989). We attribute the limited scatter of ages within a given age cluster to multiple intrusions with two batholith emplacement peaks during the Early Oligocene and Early Miocene, where magmatism might have been semi‐continuous in between these peak events, as suggested by a single Late Oligo‐
cene K‐Ar hornblende age at Balsapamba. Our zircon U‐Pb emplacement age of 25.5±0.7 Ma for the central part of the Chazo Juan‐
Telimbela batholith significantly predates pub‐
49
lished biotite ± hornblende K‐Ar ages of several batholith facies which range in age from 14.5 to 17.5 Ma and 19.1 to 21.4 Ma across the whole batholith (MMAJ/JICA 1989; MMAJ/JICA 1991; McCourt et al. 1997), again reflecting a multi‐
intrusive batholith assembly. A Re‐Os molyb‐
denite age of 19.2±0.1 Ma (Chapter 3) relates to the younger phase of intrusive activity inferred from K‐Ar data. The main pulse of plutonic activ‐
ity of the Chazo Juan‐Telimbela batholith and the youngest pulse of the Balsapamba‐Las Guardias batholith overlap in age. Likewise, the youngest plutonic pulse of Chazo Juan‐Telimbela seems to coincide with the age of the Corazon intrusion further north. Antecrystic zircons of Early Oligo‐
cene age identified in the present study could be related to an earlier phase of the Chazo Juan‐
Telimbela batholith construction, contemporane‐
ous with, or slightly postdating the first intrusive peak of Balsapamba‐Las Guardias. These age re‐
sults, in conjunction with petrographic similari‐
ties, and the spatial proximity and continuity in map view suggest that the intrusive complexes of Balsapamba‐Las Guardias and Chazo Juan‐
Telimbela, combined with the adjacent Corazon and Echeandia‐La Industria complexes and sev‐
eral smaller satellite intrusions, might represent a single large batholithic system assembled from Early Oligocene through Mid‐Miocene times. Ages reported for the Chaucha batholith range from 13.3±0.5 Ma and 13.2±0.5 Ma (K‐Ar horn‐
blende and biotite; INEMIN‐AGCD 1989), 12.8±0.6 Ma (biotite K‐Ar; Kennerley 1980), 12.0±0.6 Ma (Snelling 1970), to 9.8±0.3 (whole rock K‐Ar; Müller‐Kahle & Damon 1970). In addi‐
tion, INEMIN‐AGCD (1989) provide an age of 11±1 Ma for a dacitic porphyry intrusion. Prodeminca (2000a) note that these ages might be disturbed and could significantly postdate main batholith emplacement. Our new zircon ages confirm this notion, and provide a new ba‐
tholith minimum emplacement age of 14.84±0.06 Ma. Ubiquitous zircon antecrysts ranging in age from 15.3 to 14.8 Ma testify older pulses of magmatism, and a close‐by intrusion at the Rio Blanco prospect, dated at 15.75±0.04 Ma (zircon U‐Pb TIMS; Bineli Betsi 2007) might represent a still older, genetically related batholith facies. The K‐Ar batholith age of Müller‐Kahle & Damon (1970) is identical to our age of 9.79±0.03 for a 50 granodiorite porphyry dike intruding the batho‐
lith suggesting that the porphyry‐related hydro‐
thermal system might have thermally reset the surrounding batholith facies, but this cannot be verified due to uncertainties in the sample loca‐
tion of Müller‐Kahle & Damon (1970). Wide‐
spread porphyry‐related hydrothermal activity is documented by two Re‐Os molybdenite ages of 9.92±0.05 Ma and 9.5±0.2 Ma (Chapter 3). Age scatter in the 15‐10 Ma period might reflect mul‐
tiple intrusive events as evidenced by the occur‐
rence of an antecrystic zircon in the quartz‐
diorite porphyry intrusion dated at 10.3±0.2 Ma in the present study. The identical hornblende and biotite ages reported by INEMIN‐AGCD (1989) indicate that individual intrusive pulses were emplaced at wall rock temperatures below the K‐Ar biotite closure temperature and cooled relatively rapidly. Taken in concert, these results suggest that a main phase of the Chaucha batho‐
lith construction occurred during the Mid‐
Miocene although still earlier intrusive pulses (e.g., in the Early Miocene; Prodeminca 2000a), similar to Chazo Juan‐Telimbela, cannot be ruled out. A K‐Ar age of 16.9±0.2 Ma reported by Pratt et al. (1997) for the Paccha intrusion in the central part of the Cangrejos‐Zaruma intrusive belt has, so far, been the only available age for the whole belt. The scarcity of available data prevents further discussion of the accuracy of this K‐Ar age. In this study, a relatively tight cluster of Late Oligocene‐
Early Miocene emplacement ages was obtained from zircons for intrusions at Cangrejos, Zaruma, and Portovelo attesting coeval pluton emplace‐
ment along the whole strike length of the belt. The intrusions at Zaruma and Cangrejos contain antecrystic zircons of Mid‐Oligocene age demon‐
strating slightly older magmatic activity in this region. Combined with mainly Late Oligocene‐
Early Miocene K‐Ar and ZFT ages for volcanic rocks of the Saraguro Group in the area (Pratt et al. 1997; Hungerbühler et al. 2002, and refer‐
ences therein) this suggests the whole region un‐
derwent widespread coeval plutonism and vol‐
canism in the Late Oligocene and Early (‐Mid) Miocene. Emplacement ages of around 20 Ma obtained for the porphyry intrusions of Gaby and Papa Grande overlap within error with a 19.3±1.0 Ma K‐Ar (whole‐rock?) age for the Gaby porphyry stock reported by Prodeminca (2000a), and are further in agreement with Re‐Os molybdenite (20.6±0.1 Ma) and U‐Pb hydrothermal titanite (20.2±0.2 Ma) ages obtained for various porphyry‐related hydrothermal systems at Gaby (Chapter 3). These ages coincide with the youngest phases of batho‐
lith construction further north at Balsapamba‐Las Guardias, and testify a significant Early Miocene peak of plutonic activity in the Western Cordillera of Ecuador. Our age of 16.04±0.02 Ma for emplacement of a biotite‐bearing granodiorite porphyry dike at El Mozo, hosted by tuffs of the La Paz Formation assigned to the uppermost Saraguro Group (Prodeminca 2000b) is contemporaneous with the time range of Sta. Isabel volcanism further north (Hungerbühler et al. 2002). Furthermore, it is temporally close to the 16.9±0.2 Ma age of the Paccha intrusion (Pratt et al. 1997) and the 15.3‐
14.8 Ma zircon ages for the Chaucha batholith (this study) suggesting relatively wide‐spread post‐Saraguro magmatism in southern Ecuador. A previously reported hydrothermal alunite K‐Ar age at El Mozo is 15.4±0.7 Ma (a second age of 12.3±0.7 Ma is supposed to be inaccurate due to significant alunite concentrate contamination by barite; Prodeminca 2000b). The 15.4±0.7 Ma K‐Ar alunite age overlaps within error with our zircon age, in agreement with the general notion that porphyry intrusions and high‐sulfidation epi‐
thermal systems are genetically linked and closely correlated in time (e.g., Shinohara & Hedenquist 1997). Volcanic formations genetically related to the Quimsacocha volcanic center were not specifi‐
cally addressed by Hungerbühler et al. (2002) in their stratigraphic summary of southern Ecuador. Our age of 7.13±0.07 Ma for a Quimsacocha cal‐
dera‐hosted biotite‐hornblende‐bearing dacite dome, interpreted as emplacement age, overlaps with the time range proposed by Hungerbühler et al. (2002) and Pratt et al. (1997) for the Late Mio‐
cene, regionally widespread Tarqui Formation, and coincides on a regional scale with several shallow‐level intrusions in the area (cf. summary by Hungerbühler et al. 2002). Our new zircon U‐
Pb age further allows absolute time calibration of the volcanic stratigraphic sequence provided by Beate et al. (2001) where the early phases of the Quimsacocha volcanic center comprise andesite‐
dacite flows and breccias, followed by ignimbrite eruption and caldera formation; dacitic‐rhyolitic caldera‐hosted domes represent the final phase of activity of the volcanic center. Typical dura‐
tions of volcanic activity at long‐lived arc volca‐
noes are on the order of c. 1 m.y. (e.g., Tatara‐
San Pedro; Dungan et al. 2001). As ignimbrite ZFT ages of 5.2‐4.9 Ma and an intra‐caldera dome ZFT age of 3.6±0.3 Ma (Beate et al. 2001) significantly postdate the zircon U‐Pb age obtained for a dac‐
itic dome in the present study, the ZFT ages should either reflect exhumation, or thermal dis‐
turbance by a younger hydrothermal system. Since the Quimsacocha volcanic caldera outline is well preserved, forming a prominent positive to‐
pographic feature at about 4 km altitude, signifi‐
cant burial is not a likely option for the caldera‐
hosted facies, and hydrothermal fluids, possibly related to a blind porphyry intrusion, might have caused resetting of the ZFT system. Alternatively, these ZFT ages could be inaccurate and might not have any geological significance, or their preci‐
sion might be overestimated. The zircon U‐Pb age of 30.7±0.5 Ma for a strongly altered felsite in the Tres Chorreras prospect area assigned to the Saraguro Group (Pratt et al. 1997) places it at the base of the stratigraphic time range for the Saraguro Formation provided by Hungerbühler et al. (2002), and overlaps with the time range given by the same authors for the Loma Blanca Formation which they regard as the base of the Saraguro Group in southern Ecuador. It overlaps with the Late Eocene‐Early Oligocene ZFT ages of Saraguro Group units further north in the northern Ecuadorian arc segment (Dunkley & Gaibor 1997). Taken in combination with an age of 35.77±0.06 Ma (U‐Pb TIMS; Bineli Betsi 2007) for a voluminous quartz‐monzodiorite intrusion inferred to intrude older Saraguro Group volcan‐
ics at the close‐by Rio Blanco prospect (M. Ponce, International Minerals Corporation, pers. comm. 2009), this suggests that Early Oligocene magma‐
tism might have been more widespread in south‐
ern‐central Ecuador than previously inferred from spatially isolated outcrops of the Loma Blanca Formation, although it is clearly subordinate to Late Oligocene‐Early Miocene volcanism of the Saraguro Formation and the associated intrusive activity (Hungerbühler et al. 2002). 51
The dismembered Curiplaya porphyry intrusions hosted by the Albian Celica Formation in south‐
ernmost Ecuador were of uncertain age before this work; we therefore included them in this study of Tertiary magmatism. The porphyry intru‐
sions occur only a short distance from the com‐
posite Tangula batholith, and thus might repre‐
sent a comagmatic equivalent of this voluminous intrusive complex. A major batholith emplace‐
ment pulse is inferred at c. 110 Ma based on hornblende K‐Ar ages of 111±30 Ma for the Macará intrusion and 110±3 Ma for the Colaisaca intrusion, where an additional plagioclase K‐Ar age of 108±3 Ma is available (Kennerley 1980). Younger K‐Ar ages obtained on biotite from the same locations yielded 48±2 Ma and 93±1 Ma, respectively (Kennerley 1980). Due to their high potential for thermal disturbance the significance of these biotite ages for batholith emplacement is uncertain. Our new zircon U‐Pb age of 92.0±1.0 Ma for a Curiplaya porphyry intrusion demon‐
strates the occurrence of post‐110 Ma magma‐
tism in the area, in agreement with the age range expected from the inferred correlation of the Tangula and Peruvian Coastal batholiths (Hall & Calle 1982). although they may be of doubtful accuracy at a local scale to resolve intrusive emplacement events at higher precision. A similar conclusion can be drawn from a detailed geochronologic study conducted at the Jurassic Nambija skarn deposit in southern Ecuador: while existing K‐Ar age data on feldspar and sericite are shown to yield disturbed ages, zircon U‐Pb datations over‐
lap within error with hornblende K‐Ar ages (Chiaradia et al. 2009, and references therein). In conclusion, comparison of our U‐Pb zircon age data with existing K‐Ar ages on Tertiary intrusive rocks in Ecuador shows that several biotite and whole rock K‐Ar ages were likely disturbed by younger intrusions and their associated hydro‐
thermal systems thus failing to detect slightly older magmatic pulses (e.g., at Chaucha, Junin, Telimbela) with age offsets of 1‐4 m.y. Alterna‐
tively, our zircon ages for these intrusions were derived from previously unsampled lithologies within a given batholith; inconsistencies in pub‐
lished petrographic rock descriptions and sample locations make it difficult to evaluate this issue. In other places hornblende, biotite or whole rock K‐Ar ages coincide within error with U‐Pb zircon ages (e.g., at Gaby, Balsapamba, Cuellaje). Re‐
gionally disturbed Neogene K‐Ar age systematics in the Western Cordillera and Interandean re‐
gion, such as documented for the Late Creta‐
ceous to Early Tertiary in the Eastern Cordillera (Peltetec event of Litherland et al. 1994), are not observed. We infer that most K‐Ar ages, espe‐
cially when obtained on hornblende, can be used as a proxy for magmatism on a regional scale, Cathles, L.M., Erendi, A.H.J., and Barrie, T. (1997): How long can a hydrothermal system be sustained by a single intrusive event?: Economic Geology, v. 92, p. 766–771. 52 References Aspden, J. A., Harrison, S. H., Rundle, C. C. (1992): New geochronological control for the tectono‐magmatic evolution of the metamorphic basement, Cordillera Real, and El Oro Province of Ecuador. J S Am Earth Sc 6; 77‐96. Beate, B., Monzier, M., Spikings, R., Cotton, J., Silva, J., Bourdon, E., Eissen, J.‐P. (2001): Mio‐Pliocene adakite generation related to flat subduction in southern Ec‐
uador: the Quimasacocha volcanic center. Earth Planet Science Letters 192; 561–570. Bineli Betsi, T. (2007): The low‐sulfidation Au‐Ag de‐
posit of Rio Blanco (Ecuador): geology, mineralogy, geochronology and isotope geochemistry. Unpub‐
lished MSc. thesis, University of Geneva, 93 p. Chesley, J. (1999): Integrative geochronology of ore deposits: new insights into the duration and timing of hydrothermal circulation. In: Lambert DD, Ruiz J (eds) Application of radiogenic isotopes to ore deposit re‐
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tion in Ecuador. Mineralium Deposita 39, 204–222. Chiaradia, M., Vallance, J., Fontboté. L., Stein, H., Schaltegger, U.l, Coder, J., Richards, J., Villeneuve, M., Gendall, I. (2009): U‐Pb, Re‐Os, and 40Ar/39Ar geo‐
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phyry Cu deposits, Ecuador: implications for the Juras‐
sic metallogenic belt of the Northern Andes. Miner Deposita DOI 10.1007/s00126‐008‐0210‐6 Dungan, M. A., Wulff, A., Thompson, R. (2001): Erup‐
tive Stratigraphy of the Tatara–San Pedro Complex, 36°S, Southern Volcanic Zone, Chilean Andes: Recon‐
struction Method and Implications for Magma Evolu‐
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trology 42; 555‐626. Dunkley, P. N. & Gaibor, A. (1997): Mapa geologico de la Cordillera Occidental del Ecuador entre 2°‐3° S. es‐
cale 1/200.000. CODIGEM‐Min. Energ. Min.‐BGS publs., Quito. Foster, D. A., Kohn, B. P., Gleadow, A. J. W. (1996): Sphene and zircon fission track closure temperatures revisited: empirical calibrations from 40Ar/39Ar diffu‐
sion studies of K‐feldspar and biotite. International Workshop on Fission‐Track Dating, Ghent, Abstracts, 37. Hall, M. & Calle, J., (1982): Geochronological control for the main Tectono‐Magmatic events of Ecuador: Earth Science Review, v. 10, p. 215‐239. Henderson, W. G. (1979): Cretaceous to Eocene vol‐
canic arc activity in the Andes of northern Ecuador. Geological Society of London Journal 136; 367‐378. Hungerbühler, D., Steinmann, M., Winkler, W., Sew‐
ard, D., Egüez, A., Peterson, D. E., Helg, U., Hammer, C. (2002): Neogene stratigraphy and Andean geodynam‐
ics of southern Ecuador. Earth Science Reviews 57; 75–
124. INEMIN‐AGCD (1989): Estudio del yacimiento de cobre porfídico de Chaucha. Instituto Ec‐uatoriano de Min‐
ería, Informe final, 339 p, Quito. Kennerley J. B. (1980): Outline of the geology of Ecua‐
dor. Institute of Geological Sciences: Overseas Geology and Mineral Resources 55; 17 p. Litherland, M., Aspden, J. A., Jemielita, R. A. (1994): The metamorphic belts of Ecuador. Overseas Memoir 11. BGS, Keyworth McCourt, W.J., Duque, P., Pilatasig, L.F. and Villago‐
mez, R. 1997. Mapa geológico de la Cordillera Occi‐
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BGS publs., Quito. Prodeminca (2000a) Evaluacion de distritos mineros del Ecuador, vol 2—Depositos epitermales en la Cor‐
dillera Andina. UCP Prodeminca Proyecto MEM BIRF 36–55 EC, Quito, Ecuador. Prodeminca (2000b) Evaluacion de distritos mineros del Ecuador, vol 4—Depositos porfidicos y epi‐
mesotermales relacionados con intrusiones de las Cordilleras Occiental y Real. UCP Prodeminca Proyecto MEM BIRF 36–55 EC, Quito, Ecuador. Salazar, F. P. (2007): Geologia, alteracion y mineraliza‐
cion del porfido de cobre‐molibdeno Junin, Provincia de Imbabura, Ecuador. MSc thesis Escuela Politecnica Nacional, Quito; 72 p. Seedorff, E., Dilles, J. H., Proffett, J. M., Jr, Einaudi, M. T., Zurcher, L., Stavast, W. J. A., Johnson, D. A., Barton, M. D. (2005): Porphyry deposits: characteristics and origin of hypogene features. Econ Geol 100th Anniver‐
sary Volume, pp 251–298. Shinohara, H., Hedenquist, J. W. (1997): Constraints on Magma Degassing beneath the Far Southeast Por‐
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eral Deposits ‐ A Review of Analytical Methods, For‐
mulations, and Selected Applications. USGS Bulletin 2194. Snelling, N. (1970): K‐Ar determinations on samples from Ecuador. Int. Rep. Institute of Geo‐logical Sci‐
ences, London. Steiger, R. H. & Jäger, E. (1977): Convention on the use of decay constants in geo‐ and cosmochronology. Earth and Planetary Science Letters 36; 359‐362. Steinmann, M. (1997): The Cuenca basin of southern Ecuador: tectono‐sedimentary history and the Tertiary Andean evolution. PhD Thesis, Institute of Geology ETH Zu¨rich, Switzerland, 176 p. Van Thournout, F. (1991): Stratigraphy, magmatism and tectonism in the Ecuadorian northwestern cordil‐
lera: Metallogenic and Geodynamic implications. PhD thesis, Katholieke Universiteit Leuven, 150 p. 53
Appendix IV – Data tables Table A1: Results of U‐Pb age determinations (TIMS) Table A2: Results of U‐Pb age determinations (LA‐MC‐ICP‐MS) Table A3: Age references used for construction of Fig. 7 & Fig. 8 54 55
56 Table A2: Results of U-Pb age determinations (LA-MC-ICP-MS)
concentrations
sample
isotopic ratios
U
Th
Th/U
[ppm]
[ppm]
PS10-12
21
7
PS10-10A
26
8
PS10-13
43
16
PS10-10
26
7
0.29
PS10-20
54
34
0.63
apparent ages [Ma]
206
206
±2σ
207
±2σ
206
±2σ
corr.
206
±2σ
207
±2σ
204
207
[%]
235
[%]
238
[%]
coef.
238
0.36
348
26.7
71
0.06
72
0.0110
12
0.17
70.7
8.7
56
39
-528
1946
0.31
504
18.2
83
0.09
84
0.0123
14
0.17
78.5
11
90
73
415
1937
0.37
600
18.6
155
0.10
156
0.0134
16
0.10
85.7
13
96
143
4256
468
35.4
126
0.05
128
0.0135
22
0.17
86.2
19
52
65
357
1345
1002
20.3
45
0.09
45
0.0137
6.0
0.13
87.7
5.2
90
39
162
1060
235
207
±2σ
206
E07030
4437
PS10-5
51
19
0.37
834
24.0
95
0.08
95
0.0137
2.0
0.02
87.8
1.7
77
71
35
17
0.49
720
33.9
77
0.06
78
0.0137
16
0.21
87.8
14
55
42
-247
1214
2534
PS10-25
PS10-4
38
13
0.33
672
19.8
111
0.10
112
0.0137
6.8
0.06
87.8
5.9
93
99
220
2788
731
2457
PS10-11
41
18
0.42
474
8.0
41
0.24
48
0.0138
24
0.51
88.3
21
218
93
2038
PS10-6
128
98
0.76
1542
19.5
16
0.10
16
0.0139
2.3
0.14
89.2
2.0
95
15
249
376
PS10-17
42
27
0.63
660
20.5
57
0.09
58
0.0140
5.8
0.10
89.6
5.2
91
51
141
1373
PS10-15
43
15
0.36
810
26.0
108
0.07
109
0.0140
6.9
0.06
89.8
6.1
73
76
-449
3053
PS10-19
148
128
0.87
1740
17.4
39
0.11
40
0.0141
4.0
0.10
90.4
3.6
108
40
509
873
PS10-8
84
38
0.45
1242
20.4
22
0.10
22
0.0141
2.3
0.10
90.5
2.1
93
20
147
519
PS10-22A
58
22
0.39
1038
27.9
73
0.07
73
0.0142
2.0
0.03
90.7
1.8
69
48
-646
2058
3327
PS10-2
31
12
0.38
456
30.7
107
0.06
108
0.0144
5.4
0.05
92.2
5.0
64
66
-909
PS10-16
114
91
0.79
1536
22.0
25
0.09
25
0.0144
2.0
0.08
92.4
1.9
88
21
-35
607
PS10-22
44
13
0.30
918
19.6
64
0.10
65
0.0145
9.4
0.14
92.6
8.6
99
61
247
1514
PS10-23A
70
34
0.48
1404
25.7
57
0.08
57
0.0145
2.8
0.05
93.1
2.6
76
42
30
11
0.36
570
37.4
159
0.05
159
0.0146
7.8
0.05
93.2
7.2
53
82
-428
1530
1525
PS10-3
PS10-9
51
18
0.35
858
18.5
53
0.11
54
0.0146
3.8
0.07
93.4
3.5
105
53
370
1224
1435
6311
PS10-14
94
49
0.52
1230
17.8
63
0.11
63
0.0146
2.0
0.03
93.5
1.9
109
65
463
PS10-1T
158
145
0.92
1920
21.3
18
0.09
18
0.0147
2.0
0.11
93.9
1.9
92
16
42
423
PS10-18
49
28
0.58
702
15.4
46
0.13
46
0.0147
2.9
0.06
94.2
2.7
126
55
777
985
PS10-24
147
111
0.76
2556
22.2
32
0.09
33
0.0148
3.6
0.11
94.4
3.3
89
28
-53
795
PS10-7
53
22
0.42
636
12.1
34
0.17
34
0.0149
2.0
0.06
95.6
1.9
159
50
1257
673
PS10-21
82
32
0.39
1554
25.8
49
0.08
49
0.0150
2.6
0.05
95.7
2.5
78
37
-437
1303
PS10-23
104
94
0.90
2358
25.5
54
0.08
55
0.0152
7.5
0.14
97.1
7.3
80
42
-406
1433
PS9-5
465
189
0.41
2184
16.0
25
0.04
27
0.0042
11
0.42
26.9
3.1
36
10
696
528
PS9-16
80
42
0.53
420
7.2
73
0.08
74
0.0042
11
0.15
27.1
3.0
79
56
2206
1321
E07011
PS9-4
77
42
0.54
582
18.7
127
0.03
128
0.0043
9.6
0.08
27.5
2.6
31
40
345
3217
PS9-18
311
242
0.78
1932
19.5
26
0.03
27
0.0045
4.8
0.18
28.9
1.4
32
8
256
604
PS9-22
417
409
0.98
1992
19.7
26
0.03
26
0.0046
2.0
0.08
29.4
0.6
32
8
228
612
PS9-14
433
552
1.28
2388
19.2
16
0.03
17
0.0046
3.7
0.22
29.5
1.1
33
5
292
373
PS9-3
724
624
0.86
4392
19.9
15
0.03
16
0.0046
2.7
0.17
29.7
0.8
32
5
201
358
PS9-20
1026
1315
1.28
5742
22.6
11
0.03
11
0.0047
3.6
0.31
30.3
1.1
29
3
-98
268
PS9-2
269
238
0.88
1338
17.1
58
0.04
58
0.0047
2.9
0.05
30.3
0.9
38
22
552
1285
PS9-8
741
812
1.10
2352
16.5
13
0.04
16
0.0047
9.9
0.62
30.4
3.0
39
6
621
270
PS9-15
198
160
0.81
1182
17.2
30
0.04
30
0.0048
3.9
0.13
30.7
1.2
38
11
531
665
PS9-6A
420
300
0.71
2778
18.6
22
0.04
23
0.0048
2.0
0.09
30.7
0.6
35
8
356
509
PS9-17
1439
1348
0.94
4830
18.4
10
0.04
11
0.0048
2.0
0.19
30.8
0.6
36
4
390
234
PS9-23
1497
2013
1.34
5970
19.9
8
0.03
8
0.0048
2.5
0.31
30.9
0.8
33
3
210
180
PS9-6
190
134
0.70
462
6.1
29
0.11
30
0.0048
7.3
0.25
31.0
2.3
106
30
2510
486
PS9-13
1678
1701
1.01
10938
20.7
9
0.03
10
0.0048
3.2
0.33
31.0
1.0
32
3
115
212
PS9-24
193
120
0.62
1056
15.8
41
0.04
41
0.0049
3.3
0.08
31.5
1.0
42
17
716
873
57
Table A2 (continued)
concentrations
sample
U
Th
isotopic ratios
Th/U
206
apparent ages [Ma]
206
±2σ
207
±2σ
206
±2σ
corr.
206
±2σ
207
±2σ
235
207
±2σ
[ppm]
[ppm]
204
207
[%]
235
[%]
238
[%]
coef.
238
PS9-25
441
476
1.08
1296
12.5
57
0.05
57
0.0049
3.9
0.07
31.5
1.2
53
29
1189
206
PS9-9A
103
61
0.60
618
15.0
81
0.05
82
0.0049
6.4
0.08
31.6
2.0
45
36
825
1768
PS9-21
182
126
0.69
786
16.8
86
0.04
86
0.0050
4.1
0.05
32.0
1.3
41
34
591
1951
1142
PS9-7
609
488
0.80
3102
17.0
41
0.04
41
0.0051
2.0
0.05
32.5
0.6
41
16
560
901
PS9-10
1180
1354
1.15
6990
20.3
10
0.03
13
0.0051
7.5
0.58
32.9
2.5
35
4
160
245
PS9-19
126
134
1.06
390
4.3
32
0.17
37
0.0052
19
0.50
33.6
6.3
156
54
3055
523
PS9-12
1404
1875
1.34
6540
17.9
54
0.04
54
0.0053
2.4
0.04
33.9
0.8
40
21
442
1224
PS9-1
399
236
0.59
1356
12.7
34
0.06
34
0.0054
2.3
0.07
34.7
0.8
58
19
1160
673
PS9-9
271
150
0.55
570
5.2
77
0.15
79
0.0056
19
0.24
35.9
6.8
141
105
2769
1323
PS9-2A
133
107
0.80
390
4.5
28
0.18
30
0.0058
8.9
0.30
37.1
3.3
167
46
3011
457
PS9-11
220
173
0.79
1116
12.7
32
0.07
32
0.0066
2.0
0.06
42.3
0.8
70
21
1164
631
PS9-1A
244
86
0.35
12534
15.5
9
0.64
10
0.0716
4.4
0.44
445.7
19
500
39
759
188
PS6-14
119
61
0.51
108
58.3
463
0.01
463
0.0029
24
0.05
18.8
4.5
7
32
3374
0
PS6-13
126
59
0.47
876
16.1
865
0.03
865
0.0032
11
0.01
20.3
2.1
27
233
682
0
PS6-12
253
128
0.50
1332
21.1
48
0.02
49
0.0033
8.6
0.18
21.4
1.8
22
11
64
34
0.54
330
35.6
143
0.01
145
0.0034
26
0.18
22.0
5.6
13
19
PS6-9
78
56
0.73
444
44.8
178
0.01
178
0.0036
14
0.08
23.4
3.4
11
20
74
1364
2184
1161
PS6-4
PS6-18
217
92
0.42
1026
14.3
36
0.04
36
0.0037
5.0
0.14
23.9
1.2
36
13
922
741
PS6-18A
254
131
0.52
702
8.3
38
0.06
39
0.0039
8.9
0.23
25.1
2.2
64
24
1961
688
E06066
5208
2817
PS6-6
87
62
0.71
426
12.1
118
0.05
118
0.0040
2.0
0.02
25.6
0.5
45
52
1258
2551
PS6-2
117
51
0.43
468
10.8
60
0.05
61
0.0040
10
0.16
25.9
2.6
51
30
1484
1172
PS6-10
101
57
0.57
378
11.0
100
0.05
101
0.0040
16
0.16
25.9
4.1
50
49
1442
2049
PS6-20
160
122
0.76
372
7.5
25
0.07
26
0.0040
6.5
0.25
26.0
1.7
72
18
2136
447
PS6-17
144
174
1.21
300
5.3
34
0.11
35
0.0042
8.9
0.25
26.9
2.4
105
35
2731
560
PS6-11
238
124
0.52
660
8.5
35
0.07
35
0.0042
3.8
0.11
26.9
1.0
67
23
1930
632
PS6-19
148
50
0.34
510
6.7
53
0.09
53
0.0042
4.1
0.08
27.1
1.1
84
43
2334
924
PS6-8
97
52
0.53
408
6.0
63
0.10
66
0.0042
20
0.30
27.1
5.4
94
59
2532
1083
PS6-7
116
83
0.72
486
10.8
64
0.05
64
0.0042
3.6
0.06
27.2
1.0
53
33
1477
1237
PS6-3
92
49
0.53
288
3.1
45
0.22
51
0.0049
24
0.47
31.5
7.6
200
94
3583
3265
PS6-1
108
58
0.53
246
4.1
44
0.17
45
0.0051
12
0.26
32.5
3.8
159
66
3136
704
E07023
PS7-19
66
42
0.63
378
12.0
160
0.03
164
0.0026
37
0.23
16.5
6.1
29
48
1276
14
PS7-21
51
37
0.74
270
28.3
225
0.01
230
0.0026
46
0.20
16.7
7.7
13
29
3557
PS7-19A
57
43
0.75
354
43.4
176
0.01
185
0.0027
55
0.30
17.3
9.6
9
16
-686
2057
PS7-15
71
63
0.90
396
17.3
87
0.02
89
0.0028
17
0.19
18.3
3.1
23
20
525
1996
3372
2854
PS7-8
81
64
0.78
414
29.5
195
0.01
196
0.0029
19
0.10
18.4
3.4
13
26
PS7-13
82
51
0.63
444
52.7
366
0.01
366
0.0030
15
0.04
19.0
2.8
8
29
2786
PS7-3A
62
44
0.71
348
37.1
139
0.01
140
0.0030
21
0.15
19.1
3.9
11
16
PS7-20
80
52
0.65
570
39.1
81
0.01
82
0.0032
15
0.19
20.7
3.2
11
9
-800
2872
1502
1675
PS7-1
85
78
0.93
408
23.3
107
0.02
109
0.0032
18
0.17
20.8
3.8
19
21
-171
PS7-3
60
49
0.82
294
27.5
273
0.02
273
0.0033
8.1
0.03
21.4
1.7
17
46
-604
4008
PS7-5A
89
91
1.03
486
26.9
161
0.02
162
0.0034
13
0.08
22.0
2.8
18
28
-539
5262
0
5173
2890
PS7-10
62
47
0.76
324
11.7
42
0.04
45
0.0034
16
0.35
22.1
3.5
40
18
1326
830
PS7-20A
75
52
0.70
474
29.0
129
0.02
131
0.0034
21
0.16
22.1
4.6
16
21
4004
PS7-10A
73
66
0.91
360
44.2
140
0.01
141
0.0035
9.4
0.07
22.3
2.1
11
15
-752
2127
58 1093
Table A2 (continued)
concentrations
sample
U
Th
isotopic ratios
apparent ages [Ma]
Th/U
206
206
±2σ
207
±2σ
206
±2σ
corr.
206
±2σ
207
±2σ
235
207
±2σ
[ppm]
[ppm]
204
207
[%]
235
[%]
238
[%]
coef.
238
PS7-5
117
88
0.75
300
5.0
57
0.10
65
0.0036
33
0.50
22.9
7.5
95
60
2836
206
944
PS7-17
355
536
1.51
582
6.0
67
0.09
68
0.0038
6.2
0.09
24.6
1.5
85
55
2526
1166
PS7-6
62
50
0.81
330
24.7
122
0.02
123
0.0039
15
0.13
25.0
3.8
22
27
-317
3433
PS7-18
344
325
0.94
1230
12.9
51
0.04
53
0.0040
14
0.26
26.0
3.6
43
22
1138
1030
PS7-14
94
56
0.60
612
18.0
79
0.03
79
0.0042
8.1
0.10
26.7
2.2
32
25
1820
PS7-9
75
54
0.73
498
38.8
124
0.02
128
0.0043
32
0.25
27.6
8.7
15
20
PS7-11
139
89
0.64
780
37.2
49
0.02
49
0.0044
4.0
0.08
28.5
1.1
17
8
435
1655
1511
PS7-16A
302
416
1.38
1668
18.4
19
0.03
20
0.0045
3.5
0.18
28.8
1.0
34
6
390
433
PS7-7
178
107
0.60
1134
22.5
54
0.03
54
0.0045
4.5
0.08
29.0
1.3
28
15
-84
1340
PS7-16
142
124
0.87
786
17.9
44
0.03
44
0.0045
6.1
0.14
29.1
1.8
35
15
446
982
PS7-2
457
801
1.75
2262
18.3
24
0.04
24
0.0047
2.0
0.08
30.2
0.6
35
8
400
541
PS8-8
45
20
0.44
276
18.0
56
0.03
60
0.0033
21
0.36
21.0
4.5
25
15
437
1280
PS8-7A
73
56
0.76
408
17.6
120
0.03
123
0.0034
30
0.24
21.8
6.5
27
32
484
2908
PS8-10
53
21
0.39
384
19.4
254
0.03
254
0.0036
14
0.05
23.1
4.5
90
57
2725
1061
PS8-12
28
14
0.51
156
5.3
63
0.09
66
0.0036
19
0.30
23.1
3.1
25
64
257
2127
PS8-24
129
69
0.54
615
16.4
62
0.03
63
0.0037
5.0
0.08
24.1
1.2
31
19
640
1374
PS8-14
72
34
0.47
576
10.3
29
0.05
30
0.0038
9.5
0.31
24.4
2.3
50
15
1574
540
PS8-26
100
29
0.29
535
11.5
70
0.05
70
0.0039
5.2
0.07
24.8
1.3
46
32
1354
1401
PS8-25
136
77
0.57
670
14.3
32
0.04
34
0.0039
9.8
0.29
24.9
2.4
37
12
921
669
PS8-23
134
42
0.31
755
15.7
34
0.03
34
0.0039
2.4
0.07
25.0
0.6
34
11
734
716
PS8-15
89
40
0.45
400
6.5
43
0.08
44
0.0039
10
0.23
25.0
2.5
80
34
2379
736
PS8-18A
69
27
0.39
430
16.3
83
0.03
86
0.0039
23
0.27
25.3
5.8
33
28
652
1863
PS8-7
64
30
0.47
396
15.2
51
0.04
56
0.0039
22
0.40
25.4
5.7
36
20
794
1090
PS8-28
75
27
0.36
440
13.1
74
0.04
75
0.0040
8.3
0.11
25.5
2.1
42
31
1108
1540
PS8-5
71
48
0.68
390
14.0
67
0.04
67
0.0040
7.5
0.11
25.6
1.9
39
26
970
1396
PS8-8A
37
21
0.57
300
15.0
219
0.04
219
0.0040
20
0.09
25.9
5.1
37
79
817
993
PS8-9
59
34
0.58
612
10.9
50
0.05
52
0.0041
13
0.26
26.2
3.5
51
26
967
4677
1637
E07045
PS8-6
30
13
0.43
204
46.4
169
0.01
170
0.0041
17
0.10
26.4
4.6
12
21
1460
2327
PS8-13
55
30
0.55
402
6.2
45
0.09
49
0.0042
19
0.39
26.6
5.1
90
42
2474
778
PS8-18
77
31
0.40
475
16.0
85
0.04
85
0.0042
3.7
0.04
27.1
1.0
36
30
695
1885
PS8-19
86
48
0.56
315
4.8
74
0.12
77
0.0042
22
0.29
27.2
6.1
116
84
2882
1245
PS8-17
86
35
0.41
320
4.4
71
0.14
74
0.0043
20
0.27
28.0
5.6
131
91
3048
1188
PS8-16
80
35
0.43
315
5.5
59
0.11
60
0.0044
6.2
0.10
28.1
1.7
106
60
2684
1004
PS8-22
54
22
0.41
280
5.9
68
0.10
68
0.0044
9.0
0.13
28.2
2.5
100
65
2567
1171
PS8-20
84
36
0.43
315
6.9
54
0.09
55
0.0044
8.3
0.15
28.5
2.3
86
45
2289
946
PS8-2
56
28
0.50
300
5.2
41
0.13
42
0.0049
11
0.27
31.5
3.6
124
49
2759
674
PS8-1
94
65
0.69
288
4.6
58
0.15
59
0.0049
14
0.23
31.8
4.4
139
77
2947
956
PS8-3
49
22
0.45
282
5.3
133
0.13
139
0.0051
38
0.28
33.1
13
127
166
2719
1839
PS8-4
63
33
0.53
246
3.9
59
0.20
69
0.0059
37
0.53
37.6
14
189
120
3208
959
All errors are random errors at 2-sigma level; the additional systematic error of 1.42+/-0.54% for
for weighted mean ages presented in Table 3.
206
Pb/
238
2225
U ages is considered
59
Table A3: Age references used for construction of Fig. 7 & Fig. 8. Disturbed
ages and duplicate samples were removed.
Reference
Lithology
UTM
east
UTM
north
Aspden et al 1992
Catamayo - bt granodiorite
Aspden et al 1992
Ishpingo pluton ("unnamed")
pluton
690000
9560000
pluton
765000
9666300
Aspden et al 1992
Pichinal pluton - bt granodiorite
pluton
704500
Aspden et al 1992
age
[Ma]
±2σ
[Ma]
K-Ar bt
58
2
K-Ar bt
39
4
9599900
K-Ar bt
54
4
pluton
pluton
Aspden et al 1992
Portachuela batholith - bt-bearing felsic
porphyry
Portachuela batholith - hbl bt granodiorite
Portachuela batholith - hbl bt granodiorite
Pungala pluton - hbl bt granodiorite
677300
9472300
K-Ar bt
12
1
675500
9474400
K-Ar hbl
20
7
pluton
674500
9476500
K-Ar hbl
24
5
Aspden et al 1992
Pungala pluton - hbl bt granodiorite
pluton
768000
9800000
K-Ar hbl
42
2
pluton
768000
9796500
K-Ar bt/hbl
45
Aspden et al 1992
4
San Lucas pluton - bt granodiorite
pluton
698500
9574000
K-Ar bt
59
Aspden et al 1992
2
San Lucas pluton - hbl bt granodiorite
pluton
694800
9578500
K-Ar hbl
66
4
Aspden et al 1992
San Lucas pluton - hbl granodiorite
pluton
692800
9585700
K-Ar bt
52
2
Aspden et al 1992
pluton
693300
9584900
K-Ar bt
58
2
Barberi et al 1988
San Lucas pluton - porphyritic bt
granodiorite
Cojitambo andesite-dacite
volcanic
n/a
n/a
K-Ar
5.2
0.2
Barberi et al 1988
Mangan Fm. - dacitic lava flow
volcanic
n/a
n/a
K-Ar
8.0
0.1
Barberi et al 1988
Pisayambo Fm. - andesite
volcanic
n/a
n/a
K-Ar
12.2
0.4
Barberi et al 1988
Pisayambo Fm. - andesitic lava flow
volcanic
n/a
n/a
K-Ar
7.1
0.3
Barberi et al 1988
Pisayambo Fm. - andesitic lava flow
volcanic
n/a
n/a
K-Ar
8.1
0.1
Barberi et al 1988
Pisayambo Fm. - dacite
volcanic
n/a
n/a
K-Ar
6.1
0.6
Barberi et al 1988
Pisayambo Fm. - dacitic ignimbrite
volcanic
n/a
n/a
K-Ar
11.2
0.4
Barberi et al 1988
Pisayambo Fm. - ignimbrite
volcanic
n/a
n/a
K-Ar
15.4
0.7
Barberi et al 1988
Saraguro Fm. - andesite
volcanic
n/a
n/a
K-Ar
28.9
1.4
Barberi et al 1988
undefined andesite
volcanic
n/a
n/a
K-Ar
6.3
0.1
Beate et al 2001
Quimsacocha ignimbrites
volcanic
697400
9662500
ZFT
4.9
0.3
Beate et al 2001
Quimsacocha ignimbrites
volcanic
697400
9662500
ZFT
5.2
0.3
Bineli- Betsi 2006
Rio Blanco - microdiorite
pluton
n/a
n/a
15.8
0.04
Aspden et al 1992
Aspden et al 1992
datation
method
zircon TIMS
Bineli- Betsi 2006
Rio Blanco - qtz monzodiorite
pluton
n/a
n/a
zircon TIMS
35.8
0.06
Boland et al 1998
Apuela: Cuellaje - qtz-diorite
pluton
772702
38721
K-Ar
16.5
1.1
Boland et al 1998
Cachaco intrusion, E of Santiago
batholith
Chical-Maldonado (intrusion in San
Juan Unit) - bt-rich porphyry
pluton
789383
94048
K-Ar
34.7
1.7
pluton
811656
99597
ZFT
7.5
0.4
La Merced (Apuela satellite intrusion) qtz-diorite
Rio Naranjal gabbro intrusion
pluton
789396
71919
K-Ar
15.6
1.1
pluton
722591
38709
K-Ar
47
2
pluton
783814
95151
K-Ar
42
2
volcanic
796700
91900
ZFT
23.5
1.5
volcanic
796700
91900
ZFT
25
3
Boland et al 1998
Boland et al 1998
Boland et al 1998
Boland et al 1998
pluton
733730
0
29
3
Bourgeois et al 1990
San Eduardo intrusion, NE of Santiago
batholith
San Juan de Lachas Fm. - andesitic
breccia
San Juan de Lachas Fm. - andesitic
breccia
Santiago batholith - granodioritetonalite
Santiago batholith - granodioritetonalite
Santiago batholith - granodioritetonalite
small diorite pluton S of San Miguel de
Los Bancos
Apagua pluton - andesitic porphyry
pluton
725501
9891151
K-Ar WR
24.7
1.2
Bourgeois et al 1990
Apagua - dacite
volcanic
731254
9893145
K-Ar WR
21.3
1.1
Dunkley & Gaibor 1997
Cisarán - andesite
volcanic
730900
9777600
K-Ar
6.9
0.7
Dunkley & Gaibor 1997
Cisarán - andesite
volcanic
742900
9744100
K-Ar
7.2
0.4
Dunkley & Gaibor 1997
Molleturo diorite stock
pluton
716855
9745631
K-Ar
7.6
0.4
Dunkley & Gaibor 1997
Saraguro Group
volcanic
688700
9701200
ZFT
25.7
1.1
Dunkley & Gaibor 1997
Saraguro Group
volcanic
722800
9720800
ZFT
27.0
1.0
Dunkley & Gaibor 1997
Saraguro Group
volcanic
690700
9679300
ZFT
29.8
1.2
Dunkley & Gaibor 1997
Saraguro Group
volcanic
721700
9719200
ZFT
30.2
1.1
Dunkley & Gaibor 1997
Saraguro Group
volcanic
699900
9691800
ZFT
34.1
1.3
Dunkley & Gaibor 1997
Saraguro Group
volcanic
699800
9725200
ZFT
37.0
1.5
Dunkley & Gaibor 1997
Saraguro Group
volcanic
716400
9769200
ZFT
38.6
1.3
Boland et al 1998
Boland et al 1998
Boland et al 1998
Boland et al 1998
Boland et al 1998
Boland et al 1998
60 pluton
761551
77437
K-Ar
35.8
1.8
pluton
778256
77445
K-Ar
42
2
pluton
778246
94041
K-Ar
45
2
K-Ar
Table A3 (continued)
Reference
Lithology
UTM
east
UTM
north
Eguez 1986
La Esperie (St. Domingo) diorite
Eguez 1986
Pilaló-Zumbagua - porphyritic diorite
pluton
700326
9972354
pluton
733695
9889395
Eguez 1986
Eguez et al 1992
syntectonic intrusions in Mulaute Unit foliated diorite
Saraguro Group
pluton
733727
volcanic
Eguez et al 1992
Eguez et al 1992
Saraguro Group
Saraguro Group
Herbert & Pichler 1983
datation
method
age
[Ma]
±2σ
[Ma]
K-Ar WR
38.6
1.9
K-Ar WR
24.7
1.2
9966819
K-Ar hbl
48.3
0.6
728500
9762000
K-Ar plag
21.0
1.0
volcanic
728500
9764800
K-Ar plag
27.0
0.9
volcanic
724000
9746700
K-Ar
35.9
0.9
Amaluza pluton - granodiorite
pluton
792500
9712314
K-Ar bt
34
1
Herbert & Pichler 1983
San Lucas pluton - bt granite
pluton
694355
9588998
K-Ar bt
52
2
Hughes et al 1998
Chaupicruz - granodiorite
pluton
717002
9900465
ZFT
7.0
0.3
Hughes et al 1998
pluton
711443
9917057
K-Ar hbl
38.1
0.4
Hughes et al 1998
El Tigre, R. Hugshatambo - granodiorite
R. Quindigua - granodiorite
pluton
728141
9911519
K-Ar bt/hbl
14.8
0.14
Hughes et al 1998
Zumbagua - porphyritic tonalite
volcanic
733698
9894925
K-Ar bt/hbl
6.3
0.7
Hungerbühler et al 2002
Chinchin Fm.
volcanic
739467
9680826
ZFT
43
4
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
678920
9562364
ZFT
25
3
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
685992
9538147
ZFT
27
4
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
680590
9544440
ZFT
29
3
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
646676
9558934
ZFT
31
3
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
685900
9538200
ZFT
33
3
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
687184
9537098
ZFT
33
4
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
699139
9546868
ZFT
36
7
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
700498
9519437
ZFT
41
5
Hungerbühler et al 2002
Loma Blanca Fm.
volcanic
647582
9559391
ZFT
42
3
Hungerbühler et al 2002
Rodanejo pluton
pluton
672059
9546653
ZFT
39
6
Hungerbühler et al 2002
Sacapalca Fm.
volcanic
646051
9555116
ZFT
67
6
Hungerbühler et al 2002
Saraguro Fm.
volcanic
10365
24540
ZFT
19
6
Hungerbühler et al 2002
Saraguro Fm.
volcanic
9820
25430
ZFT
19
4
Hungerbühler et al 2002
Saraguro Fm.
volcanic
681315
9630791
ZFT
19.1
1.4
Hungerbühler et al 2002
Saraguro Fm.
volcanic
10450
27840
ZFT
20
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
735092
9706274
ZFT
21
2
Hungerbühler et al 2002
Saraguro Fm.
volcanic
671534
9632381
ZFT
21
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
735732
9702400
ZFT
21.2
1.6
Hungerbühler et al 2002
Saraguro Fm.
volcanic
695744
9693134
ZFT
23
2
Hungerbühler et al 2002
Saraguro Fm.
volcanic
13480
31070
ZFT
23
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
697035
9625187
ZFT
23
2
Hungerbühler et al 2002
Saraguro Fm.
volcanic
735533
9699850
ZFT
23.2
1.8
Hungerbühler et al 2002
Saraguro Fm.
volcanic
697311
9624303
ZFT
23
2
Hungerbühler et al 2002
Saraguro Fm.
volcanic
726472
9673198
ZFT
26
2
Hungerbühler et al 2002
Saraguro Fm.
volcanic
730250
9675252
ZFT
26.0
1.8
Hungerbühler et al 2002
Saraguro Fm.
volcanic
680200
9630300
ZFT
26
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
714257
9685269
ZFT
26
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
12250
29500
ZFT
26
5
Hungerbühler et al 2002
Saraguro Fm.
volcanic
698991
9674203
ZFT
27
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
730307
9675184
ZFT
27
4
Hungerbühler et al 2002
Saraguro Fm.
volcanic
730062
9675245
ZFT
27
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
729721
9675308
ZFT
28
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
720997
9667853
ZFT
28
3
Hungerbühler et al 2002
Saraguro Fm.
volcanic
725302
9693138
ZFT
29
3
Hungerbühler et al 2002
St. Isabel Fm.
volcanic
698813
9656909
ZFT
8
2
Hungerbühler et al 2002
St. Isabel Fm.
volcanic
699931
9642136
ZFT
15.9
1.6
Hungerbühler et al 2002
St. Isabel Fm.
volcanic
688300
9629400
ZFT
18
2
Hungerbühler et al 2002
St. Isabel Fm.
volcanic
689714
9631727
ZFT
18
3
Hungerbühler et al 2002
St. Isabel Fm.
volcanic
683296
9633982
ZFT
18.4
1.6
Hungerbühler et al 2002
St. Isabel Fm.
volcanic
698428
9640804
ZFT
19
2
Hungerbühler et al 2002
Tarqui Fm.
volcanic
732155
9672200
ZFT
5.1
0.6
Hungerbühler et al 2002
Tarqui Fm.
volcanic
729094
9687475
ZFT
5.5
0.6
Hungerbühler et al 2002
Tarqui Fm.
volcanic
731940
9672631
ZFT
5.8
0.8
Hungerbühler et al 2002
Tarqui Fm.
volcanic
728773
9664281
ZFT
6.0
1.0
Hungerbühler et al 2002
Tarqui Fm.
volcanic
720818
9672195
ZFT
6.1
1.0
Hungerbühler et al 2002
Tarqui Fm.
volcanic
729453
9664218
ZFT
6.3
0.8
Hungerbühler et al 2002
Tarqui Fm.
volcanic
13550
28200
ZFT
6.3
1.0
61
Table A3 (continued)
Reference
Lithology
Hungerbühler et al
2002
Hungerbühler et al
2002
Hungerbühler et al
2002
INEMIN-AGCD 1989
Tarqui Fm.
Tarqui Fm.
Tarqui Fm.
volcanic
723285
9725220
ZFT
6.8
0.8
Chaucha batholith - tonalite
pluton
666734
9690407
K-Ar bt
13.2
0.5
INEMIN-AGCD 1989
Chaucha batholith - tonalite
pluton
666734
9690407
K-Ar hbl
13.3
0.5
INEMIN-AGCD 1989
Chaucha dacitic porphyry
pluton
666734
9690407
K-Ar bt
11.0
1.0
Jaillard et al 1996
Palo Blanco pluton - granodiorite
pluton
669200
9541800
K-Ar plag
26.6
1.6
Kennerley 1980
Amaluza pluton - granodiorite
pluton
770575
9712820
K-Ar hbl
Kennerley 1980
andesite; Hungerbühler (1997) assigns to
El Descanso intrusion
andesite; Hungerbühler (1997) assigns to
El Descanso intrusion
andesite; Hungerbühler (1997) assigns to
St. Isabel Fm.
andesitic porphyry; Hungerbühler (1997)
assigns to St. Isabel Fm.
pluton
736041
9686336
pluton
736690
volcanic
volcanic
Kennerley 1980
Kennerley 1980
Kennerley 1980
UTM
east
UTM
north
volcanic
728865
9663789
volcanic
744998
9700442
datation
method
age
[Ma]
±2σ
[Ma]
ZFT
6.6
0.8
ZFT
6.7
0.8
49
2
K-Ar WR
19.7
0.5
9686612
K-Ar WR
21
0.6
673135
9632347
K-Ar WR
19.5
0.4
698160
9643179
K-Ar WR
14.2
0.5
0.6
Kennerley 1980
Chaucha batholith - granodiorite
pluton
675428
9679338
K-Ar bt
12.8
Kennerley 1980
El Tingo pluton - granodiorite
pluton
678559
9558619
K-Ar bt
50
3
Kennerley 1980
Las Guardias pluton - qtz-diorite
pluton
711358
9800935
K-Ar hbl
20
3
Kennerley 1980
Portachuela batholith - granite
pluton
674723
9493201
K-Ar bt
29.0
0.8
Kennerley 1980
volcanic
700010
9641978
K-Ar WR
21.4
0.8
Kennerley 1980
rhyolite; Hungerbühler (1997) assigns to St.
Isabel Fm.
San Lucas pluton - granodiorite
pluton
702191
9573316
K-Ar bt
Kennerley 1980
Saraguro Group - rhyolite
volcanic
672821
9632716
Lavenu et al 1992
Biblian Fm - rhyolitic tuff
volcanic
735300
Lavenu et al 1992
pluton
735300
Lavenu et al 1992
Cojitambo - andesite; Hungerbühler (1997):
this Cojitambo sample is intrusive, whereas
younger Cojitambo ages are from extrusive
rocks
Mangan Fm. - rhyolitic tuff
volcanic
Lavenu et al 1992
Pisayambo Fm - andesite
volcanic
Lavenu et al 1992
Pisayambo Fm - andesite
volcanic
694400
Lavenu et al 1992
Pisayambo Fm - andesite
volcanic
737400
9893100
K-Ar plag
9.1
0.5
Lavenu et al 1992
Saraguro Fm - andesite
volcanic
728000
9764100
K-Ar WR
21.0
1.0
Lavenu et al 1992
Saraguro Fm - andesite
volcanic
724300
9747500
K-Ar plag
35.5
1.3
Lavenu et al 1992
Saraguro Fm. - andesite
volcanic
733500
9686600
K-Ar plag
35.3
0.9
McCourt et al 1997
Balsapamba - tonalite-granodiorite
pluton
700235
9806476
K-Ar bt/hbl
33.1
0.4
McCourt et al 1997
Chaso Juan - tonalite-granodiorite
pluton
708060
9845175
K-Ar bt/hbl
19.5
0.3
McCourt et al 1997
Chaso Juan - tonalite-granodiorite
pluton
708060
9845175
K-Ar bt/hbl
20.7
0.2
McCourt et al 1997
Corazon batholith - tonalite-granodiorite
pluton
722548
9868388
K-Ar bt/hbl
14.1
0.3
McCourt et al 1997
Corazon batholith - tonalite-granodiorite
pluton
728113
9867278
K-Ar bt/hbl
14.8
0.2
McCourt et al 1997
Corazon batholith - tonalite-granodiorite
pluton
728117
9872808
K-Ar bt/hbl
14.8
0.4
McCourt et al 1997
Corazon batholith - tonalite-granodiorite
pluton
728117
9872808
K-Ar bt/hbl
16.1
0.2
McCourt et al 1997
La Industria tonalite-granodiorite
pluton
694671
9806481
K-Ar bt/hbl
23.1
0.8
McCourt et al 1997
La Industria tonalite-granodiorite
pluton
694676
9812010
K-Ar bt/hbl
25.6
0.3
McCourt et al 1997
La Industria tonalite-granodiorite
pluton
689108
9806486
K-Ar bt/hbl
26.5
0.7
McCourt et al 1997
Las Guardias tonalite-granodiorite
pluton
705788
9795411
K-Ar bt/hbl
33.4
0.3
McCourt et al 1997
Las Guardias tonalite-granodiorite
pluton
705788
9795411
K-Ar bt/hbl
34.3
0.8
McCourt et al 1997
stock intruding Yunguilla Unit at Juan de
Velasco - porphyritic granodiorite
pluton
733615
9800911
K-Ar
10.1
0.2
McCourt et al 1997
Tambana pluton
pluton
705776
9784352
K-Ar bt/hbl
25.4
0.2
McCourt et al 1997
Telimbela - tonalite-granodiorite
pluton
705811
9818635
K-Ar bt/hbl
19.1
0.8
McCourt et al 1997
Telimbela - tonalite-granodiorite
pluton
700246
9817534
K-Ar bt/hbl
20.0
0.4
McCourt et al 1997
Telimbela - tonalite-granodiorite
pluton
700246
9817534
K-Ar bt/hbl
21.4
0.2
McCourt et al 1997
tonalite dike cutting Apagua Formation
pluton
728125
9883869
K-Ar bt/hbl
23.7
0.5
MMAJ/JICA 1989
Balsapamba - qtz-diorite
pluton
707560
9807830
K-Ar hbl
25.7
0.9
MMAJ/JICA 1989
Chaso Juan - granodiorite
pluton
706170
9845140
K-Ar bt
20.9
0.7
MMAJ/JICA 1989
La Industria - qtz-diorite
pluton
690970
9825260
K-Ar hbl
25.5
0.9
MMAJ/JICA 1989
Telimbela - qtz-diorite
pluton
703680
9816010
K-Ar bt
19.4
0.6
MMAJ/JICA 1989
Las Guardias - qtz-diorite
pluton
708140
9798660
K-Ar hbl
30.1
1.1
MMAJ/JICA 1991
Chaso Juan - diorite
pluton
n/a
n/a
K-Ar
17.5
0.6
MMAJ/JICA 1991
Telimbela - hbl qtz-diorite
pluton
n/a
n/a
K-Ar
15
3
MMAJ/JICA 1991
Telimbela - qtz porphyry
pluton
n/a
n/a
K-Ar
15.7
1.0
62 63
1
K-Ar WR
26.8
0.7
9701400
K-Ar plag
22.0
0.8
9695800
K-Ar plag
7.1
0.3
733500
9697700
K-Ar plag
16.3
0.7
739100
9760300
K-Ar WR
7.9
0.4
9596400
K-Ar plag
8.2
0.4
Table A3 (continued)
Reference
Lithology
UTM
east
UTM
north
MMAJ/JICA 1992
Apuela: Cuellaje - granodiorite
MMAJ/JICA 1992
Apuela: Junin - granodiorite
pluton
778270
42042
pluton
767136
27657
MMAJ/JICA 1992
Cuellaje andesite porphyry
pluton
778270
MMAJ/JICA 1992
Cuellaje qtz porphyry
pluton
MMAJ/JICA 1992
Junin diorite porphyry
MMAJ/JICA 1992
Junin qtz porphyry
Müller-Kahle &
Damen 1970
OLADE 1980
Pichler & Aly 1983
datation
method
age
[Ma]
±2σ
[Ma]
K-Ar bt
13.0
0.6
K-Ar hbl
14.5
0.2
42042
K-Ar WR
11.1
0.6
778270
42042
K-Ar WR
8.8
0.4
pluton
767136
27657
K-Ar WR
7.3
0.3
pluton
767136
27657
K-Ar WR
6.1
0.2
Chaucha batholith
pluton
675053
9676574
K-Ar WR?
9.8
0.3
Cojitambo andesite-dacite
volcanic
735316
9695831
K-Ar
6.3
0.2
Pungala pluton - granodiorite
pluton
770721
9802708
K-Ar bt
41.3
1.6
Pratt et al 1997
NE Uzhcurrumi - qtz-diorite
volcanic
661100
9635129
K-Ar
19.9
0.2
Pratt et al 1997
Paccha granitoid
pluton
644386
9601984
K-Ar
16.9
0.2
Pratt et al 1997
Saraguro Fm.
volcanic
650200
9591600
ZFT
21.5
1.6
Pratt et al 1997
Saraguro Fm.
volcanic
632400
9635400
ZFT
23.2
1.6
Pratt et al 1997
Saraguro Fm.
volcanic
661900
9650900
ZFT
28
2
Pratt et al 1997
Saraguro Fm. - dacitic tuff
volcanic
690200
9629700
ZFT
22
2
Pratt et al 1997
Saraguro Fm. - dacitic tuff
volcanic
690700
9629300
ZFT
27
2
Pratt et al 1997
Saraguro Fm. - ignimbrite
volcanic
702934
9614065
ZFT
22.5
1.8
Pratt et al 1997
Saraguro Fm. - ignimbrite
volcanic
703862
9618808
ZFT
25.0
1.8
Pratt et al 1997
Shagli intrusion - granodiorite
pluton
683353
9651682
K-Ar
17.6
0.6
Pratt et al 1997
Tarqui Fm. - dacite lava flow
volcanic
694100
9595400
ZFT
9.6
1.0
Pratt et al 1997
undefined intrusion
pluton
658700
9651500
ZFT
13.9
1.0
Prodeminca 2000
Apuela batholith: granodiorite
pluton
n/a
n/a
K-Ar hbl
16.5
0.8
Prodeminca 2000
Junin - qtz porphyry
pluton
n/a
n/a
K-Ar bt/hbl
5.9
0.1
Prodeminca 2000
Junin - qtz porphyry
pluton
n/a
n/a
K-Ar bt/hbl
7.9
0.3
Rivera et al 1992
Saraguro Fm. - ignimbrite
volcanic
697700
9675200
K-Ar bt
26.0
0.8
Rivera et al 1992
Saraguro Fm. - ignimbrite
volcanic
696300
9676200
K-Ar bt
27.0
0.7
Snelling 1970
Chaucha batholith - granodiorite-tonalite
pluton
666734
9690407
K-Ar
12.0
0.6
Spikings et al. 2005
Saraguro Group volcanic rock
volcanic
698100
9724200
ZFT
36
3
Steinmann 1997
Calera pluton - granite
pluton
650061
9591480
ZFT
26.5
1.8
Steinmann 1997
Cisarán - andesite
volcanic
725300
9693000
ZFT
6.8
0.8
Steinmann 1997
El Prado pluton - granite
pluton
658870
9578475
ZFT
24
2
Steinmann 1997
Porotillos pluton - granite
pluton
653857
9632490
ZFT
20
4
Steinmann 1997
San Antonio pluton - granite
pluton
662162
9634470
ZFT
20
3
this study
Apuela: Cuellaje granodiorite (E060206)
pluton
772701
44253
zircon TIMS
12.87
0.08
this study
Balsapamba granodiorite (E06140)
pluton
708028
9809786
zircon TIMS
21.46
0.08
this study
pluton
708028
9809786
zircon TIMS
21.22
0.17
this study
Balsapamba granodiorite porphyry
(E06131)
Cangrejos qtz-diorite (E06066)
pluton
633163
9614248
25.7
1.0
this study
Chaucha dacitic porphyry (Tunas; E07005)
pluton
675055
9677495
zircon
ICPMS
zircon TIMS
9.79
0.03
this study
Chaucha granodiorite batholith (E07003)
pluton
675055
9677495
zircon TIMS
14.84
0.06
this study
El Mozo granodiorite porphyry (E07018)
pluton
714415
9618449
zircon TIMS
16.04
0.02
this study
Gaby plag-hbl porphyry (E05083)
pluton
644460
9657265
zircon TIMS
20.26
0.06
this study
Junin granodiorite porphyry (E07032)
pluton
755998
33186
zircon TIMS
9.01
0.05
this study
Papa Grande plag-hbl porphyry (E05090)
pluton
645570
9656158
zircon TIMS
19.89
0.06
this study
Portovelo plag-hbl porphyry (E06112)
pluton
653987
9589516
zircon TIMS
24.04
0.06
this study
Quimsacocha altered dacite dome
(E06017)
Saraguro at Tres Chorreras - felsite
(E07011)
Telimbela granodiorite (E07045)
pluton
697346
9662472
zircon TIMS
7.13
0.07
volcanic
663344
9649685
pluton
714543
9834663
zircon
ICPMS
zircon
ICPMS
zircon
ICPMS
Ar-Ar gm
this study
this study
this study
Zaruma granodiorite (E07023)
pluton
653270
9599944
Vallejo 2007
Cizaran Fm. - andesite
volcanic
740461
9806100
30.7
0.6
25.5
0.7
20.9
1.1
12.2
2.2
Vallejo 2007
Macuchi Unit - andesite
volcanic
784106
95939
Vallejo 2007
Macuchi Unit - andesite
volcanic
725129
9965028
Ar-Ar gm
35.1
1.7
Ar-Ar plag
42.6
1.3
Vallejo 2007
Pilalo Fm. - andesite
volcanic
771610
9996629
Ar-Ar px
64.3
0.4
Vallejo 2007
Pilalo Fm. "intrusion" - andesite
volcanic
733838
9919774
Ar-Ar hbl
34.8
1.4
Vallejo 2007
Rio Cala Unit - basaltic andesite
volcanic
787170
27797
Ar-Ar px
Vallejo 2007
San Juan de Lachas Fm. - andesite
volcanic
806395
83179
Ar-Ar
Vallejo 2007
Silante Fm. - basalt
volcanic
763379
9995871
Vallejo 2007
Silante Fm. - andesite
volcanic
766935
2688
Vallejo 2007
Silante Fm. - andesite
volcanic
768285
Vallejo 2007
Silante Fm. - basalt
volcanic
763379
67
7
32.9
1.2
Ar-Ar gm
66
2
Ar-Ar gm
58
2
1600
Ar-Ar gm
61.0
1.1
9995871
Ar-Ar gm
66
2
63
Table A3 (continued)
Reference
Lithology
van Thournout 1991
Apuela pluton - granodiorite
pluton
776600
30100
K-Ar bt
15.8
0.6
van Thournout 1991
Apuela pluton - granodiorite
pluton
776600
30100
K-Ar hbl
18.5
0.9
van Thournout 1991
pluton
n/a
n/a
K-Ar
20
3
pluton
n/a
n/a
K-Ar
36
2
van Thournout 1991
hbl-rich intrusion hosted by San Juan de Lachas
Fm.
hbl-rich intrusion hosted by San Juan de Lachas
Fm.
Macuchi Unit. - gabbro
pluton
791000
91500
K-Ar hbl
45
9
van Thournout 1991
Maldonado pluton - granodiorite
pluton
822200
101500
K-Ar bt
8.9
0.4
van Thournout 1991
Rio Babosa granodiorite
pluton
784100
98500
K-Ar hbl
40
3
van Thournout 1991
Tandapi Unit - diorite
pluton
806400
83600
K-Ar hbl
32.6
1.3
van Thournout 1991
UTM
east
UTM
north
datation
method
age
[Ma]
±2σ
[Ma]
Acronyms: qtz - quartz, plag - plagioclase, px - pyroxene, hbl - honrblende, gm - groundmass, WR - whole rock, ZFT - zircon fission track
References
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Eguez, A., Dugas, F., Bonhomme, M. (1992): Las Unidades Huigra y Alausi en la Evolucion Geodinamica del Valle Interandino del Ecuador. Boletin Geologico Ecuatoriano 3; 4756.
Herbert, H. J. & Pichler, H. (1983): K-Ar ages of rocks from the Eastern Cordillera of Ecuador. Zeitschrift der Deutschen Geologischen Gesellschaft 134; 483-493.
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64 CHAPTER III THE MIOCENE METALLOGENIC BELT OF ECUADOR: CONSTRAINTS FROM NEW Re‐Os MOLYBDENITE AND U‐Pb TITANITE AGES OF PORPHYRY‐RELATED ORE DEPOSITS Abstract This study presents Re‐Os molybdenite and U‐Pb titanite ages related to hydrothermal pulses of mineraliza‐
tion and alteration of latest Oligocene and Miocene porphyry‐related ore deposits in Ecuador. Molybdenite associated with potassic‐phyllic alteration at the Junin Cu‐Mo porphyry deposit yielded ages of 6.63±0.04 Ma and 6.13±0.03 Ma. Re‐Os ages of molybdenite associated with potassic alteration at the Telimbela and Balsapamba porphyry systems are 19.2±0.1 Ma and 21.5±0.1 Ma, respectively. At the Chaucha Cu‐Mo por‐
phyry system, Re‐Os ages of 9.92±0.05 Ma (Tunas‐Naranjos) and 9.5±0.2 Ma (Gur‐Gur) were obtained for molybdenite associated with potassic‐phyllic alteration. At the Gaby Au‐Cu porphyry, a Re‐Os molybdenite age of 20.6±0.1 Ma for a sulfide‐cemented hydrothermal breccia (possibly related to phyllic alteration), and a U‐Pb age of 20.17±0.16 Ma for titanite associated with Na‐Ca alteration were obtained. At the Tres Chor‐
reras polymetallic deposit, Re‐Os molybdenite ages are 12.93±0.07 Ma and 12.75±0.07 Ma, and are associ‐
ated with an intrusion‐related hydrothermal breccia and a polymetallic vein, respectively. Molybdenite as‐
sociated with Na‐Ca alteration at the Cangrejos Au‐Cu porphyry system yielded an age of 23.5±0.1 Ma. Our new geochronologic data allow us to infer that the Miocene metallogenic belt of northern‐central Peru extends northwards into southern Ecuador, and potentially further north until Colombia. Intersections of the Andean (NNE‐) trending magmatic arc with arc‐transverse faults and lineaments related to suture zone geometries and block rotation in southern Ecuador represent highly prospective sites for Miocene por‐
phyry‐related mineralization. Porphyry‐related ore deposits in Ecuador are often associated with intrusive clusters of batholith dimen‐
sions, where porphyry‐related pulses of hydrothermal activity often occur towards the end of batholith as‐
sembly. Thus, while batholith complexes may mark structurally favorable sites for mineralization, extensive shallow crustal magmatism during peak periods of batholith construction may be disadvantageous for the formation and preservation of porphyry‐related ore deposits. The lack of Quaternary arc volcanic cover sequences due to a local flat slab setting, and overall favorable erosion levels are key parameters to preserve and expose widespread Miocene epithermal and porphyry Cu mineralization in southern Ecuador. In the Western Cordillera of Ecuador porphyry Cu mineralization has locally been preserved, whereas the deeply eroded cores of porphyry systems are exposed at other loca‐
tions where significant parts of the mineralization have been removed. Although possibly applicable for single ore deposits, a general, direct spatio‐temporal association between Miocene ore deposit formation and seamount chain ("ridge") subduction or regional compressive pulses, as sometimes proposed for parts of the central and southern Andes, is not observed in Ecuador. 65
Introduction Porphyry‐related (porphyry Cu, epithermal, and Cordilleran vein type) ore deposits in western South America typically occur in elongated metal‐
logenic belts of several 100 km strike length which regionally link deposits of a similar age (Sil‐
litoe 1988). Additional sub‐belts may be defined where the density of ore deposits and geochro‐
nologic control on the timing of mineralization are sufficiently high (e.g., central‐northern Peru; Noble & McKee 1999). Porphyry‐related ore de‐
posits of the Andes are intimately associated with arc magmatism resulting from the subduction of the Farallon/Nazca plate at the South American margin (Sillitoe 1988). Consequently, metal‐
logenic belts tend to follow the overall spatio‐
temporal distribution of arc magmatism which is mainly dictated by the subducting slab geometry and upper plate structures (e.g., Kay et al. 1999; Tosdal & Richards 2001; Richards 2003). Further geodynamic and tectonomagmatic con‐
trols may operate, causing particularly prolific intervals of mineralization in certain belt seg‐
ments, where it is important to distinguish in‐
creased rates of porphyry‐related ore deposit formation from optimum conditions of ore de‐
posit exposure and preservation (at constant rates of deposit formation; Wilkinson & Kesler 2009). Possible factors which have been pro‐
posed to show a positive feedback with porphyry‐
related mineralization at regional to local scales include the subduction of bathymetric anomalies such as seamount chains ("ridges"), and their ef‐
fects on crustal deformation (e.g., Rosenbaum et al. 2005; Cooke et al. 2005), intense hydration of the crust and evolving arc magmas by flat slab dehydration (James & Sacks 1999) or amphibole break‐down (Kay et al. 1999), and a broadly fa‐
vorable stress regime and its bearing on the ge‐
ometry of crustal structures and transcrustal magma ascent (Tosdal & Richards 2001; Richards 2003). Furthermore, progressive volatile enrich‐
ment of porphyry intrusive parental melts at mid‐ to deep crustal levels might represent a favorable magmatic preconditioning stage for subsequent intrusion‐related mineralization at shallower lev‐
els (Rohrlach & Loucks 2005; Chiaradia et al. 2009a). 66 While the broad extent of the Miocene metal‐
logenic belts of Peru and Chile is relatively well established, the northward belt continuation into Ecuador is less certain. In fact, in his seminal in‐
vestigation of intrusion‐related metallogenic belts of the Andes, Sillitoe (1988) depicted a pro‐
nounced metallogenic gap for Ecuador as only a single Miocene porphyry Cu deposit (Chaucha) was relatively well known at that time. Over the last two decades, Ecuador has increasingly moved into the focus of exploration activities of public and privately owned companies as well as several government agencies, resulting in the discovery and re‐assessment of a significant number of Miocene and older porphyry‐related ore deposits (Prodeminca 2000a, 2000b; USGS 2009). In combination with their Jurassic equiva‐
lents, the Miocene ore deposits contain the bulk of the country's resources in Cu, Mo, Au, and Ag (Prodeminca 2000a). Miocene mineralization types mainly comprise porphyry‐style (e.g., Junin, Chaucha, Gaby), high sulfidation (e.g., Quimsaco‐
cha), intermediate sulfidation (e.g., Portovelo‐
Zaruma), and low sulfidation (e.g., Rio Blanco) epithermal deposits. Although some of these ore deposits have a long‐standing history of artisanal production, partly since Inca times, many of Ec‐
uador's mineral resources remain undeveloped to date (Prodeminca 2000a, 2000b; Spencer et al. 2002; USGS 2009). Where quantitatively assessed (e.g., Tab. 1, for deposits investigated in this study), their tonnage seems to lag behind that of some giant ore deposits of the Miocene Peruvian metallogenic belt (e.g., Noble & McKee 1999; Rosenbaum et al. 2005). In this contribution we present ten new Re‐Os molybdenite ages related to porphyry‐style or epithermal mineralization (complemented by a U‐Pb titanite age related to hydrothermal altera‐
tion) at several Miocene ore deposits in southern, central, and northern Ecuador (Figs. 1, 2). Com‐
bined with recent geochronologic works on Late Tertiary igneous rocks (Chapter 2) and available literature data (Prodeminca 2000a, 2000b, and references therein), our new geochronologic data allow us to assess the Miocene metallogenic po‐
tential of Ecuador, its connectivity with the northern‐central Peruvian Miocene metallogenic belt, and its relation with the geodynamic evolu‐
tion of the Ecuadorian margin. Figure 1: Simplified geological map of Tertiary arc magmatic units at the NW South American margin, location of irregular
bathymetric features of the sub‐
ducting Nazca plate, and Late
Oligocene‐Miocene intrusion‐
related ore deposits. Only ore
deposits dated in this study are
displayed for Ecuador; deposit data for northern Peru and
southern Colombia from Sillitoe
(1988) and Noble et al. (2004),
slightly modified to account for recent discoveries. Major struc‐
tures (undifferentiated; mostly thrust faults which have been
variably reactivated during the
Tertiary) from references com‐
piled in Chapter 1, and addition‐
ally Mégard (1984) and McNulty
et al. (1998) for northern Peru;
major structures of the Eastern Cordilleras of Ecuador and Peru are not shown. Ore deposits in
Peru all plot in the Mid‐Miocene to Early Pliocene Cu metal‐
logenic belt of Sillitoe (1988); ore
deposits in southern Ecuador fit
well into the northward projec‐
tion of this belt in space and time. Ore deposit density further
north is lower, but suggests that the Miocene Cu belt might be
broadly continuous into Colom‐
bia, following the magmatic arc.
Hypothetic positions of older,
already subducted oceanic fea‐
tures such as the Inca plateau are not shown, but might be of
metallogenetic significance (e.g.,
Rosenbaum et al. 2005). The Curiplaya porphyry intrusions in
southernmost Ecuador are of Late Cretaceous age (Chapter 1)
and do not form part of the Mio‐
cene metallogenic belt. 67
Regional geology and geody‐
namic setting Since the Late Cretaceous the Ecuadorian sub‐
duction system has been influenced by a series of major geodynamic events starting with the c. 75‐
70 Ma accretion of oceanic plateau fragments which floor the present day forearc region and possibly parts of the Interandean Depression (e.g., Vallejo et al. 2009). Throughout the Tertiary oblique plate convergence between the Faral‐
lon/Nazca and the South American plates has been accommodated by combined oblique sub‐
duction slip and trench‐parallel forearc sliver dis‐
placement where the former is the dominant mechanism at the present day (Daly 1989; Ego et al. 1996). The offshore Grijalvas scarp separates Farallon and Nazca seafloor, the latter progres‐
sively produced since the break‐up of the Faral‐
lon plate in the Early Miocene (Lonsdale 2005). The ENE‐trending scarp currently intersects the Ecuadorian trench at 3°S implying that old Faral‐
lon crust is now subducted below southern Ecua‐
dor (and further south in Peru) whereas young Nazca crust is subducted below central‐northern Ecuador (and further north in Colombia); Mio‐
cene plate motions (Somoza 1998) dictate that the scarp progressively swept southwards along the margin (Gutscher et al. 1999). The Carnegie Ridge seamount chain collided with the Ecuador‐
ian margin in the Late Miocene (‐Pliocene?) al‐
though the exact timing of initial collision is still a matter of debate, and complicated by jumps of the Cocos‐Nazca spreading center and a possibly segmented seamount track (Lonsdale & Klitgord 1978; Daly 1989; Gutscher et al. 1999; Figure 2: Geological map of the southern Ecuadorian Sierra region showing position of mineral deposits investigated in
this study (Ar‐Ar data is still pending and will be supplemented as it becomes available). Note that a larger number of ore deposits occurs in this area (Prodeminca 2000a, b). Black diamonds correspond to U‐Pb zircon intrusive ages (in Ma: Chapter 1, and Bineli Betsi, 2007, for Rio Blanco intrusions). White diamonds correspond to intrusive ages obtained
by K‐Ar (and in one case zircon fission track) geochronology (Aspden et al. 1992; Pratt et al. 1997). Only ages consid‐
ered as relevant for intrusion emplacement are shown. Ages of Saraguro Group volcanics are mostly 19‐29 Ma (ZFT; Hungerbühler et al. 2002). Adapted from Litherland et al. (1994), Pratt et al. (1997) and Dunkley & Gaibor (1997).
68 Spikings et al. 2001; Witt et al. 2006). Ridge colli‐
sion seems to have caused shallowing of the sub‐
duction angle from c. 30‐35° to 25‐30° below cen‐
tral‐northern Ecuador in the Late Miocene‐
Pliocene, whereas the subduction angle had been broadly constant during the Oligocene‐Miocene (Guillier et al. 2001; Chapter 2). Establishment of a flat slab geometry below northern Peru and southernmost Ecuador, associated with a gap in arc magmatism, initiated in the Mid‐ to Late Mio‐
cene (e.g., James & Sacks 1999; Gutscher et al. 1999; Chapter 2). The flat slab segment and the accompanying cessation of arc magmatism seem to have broadened progressively towards south‐
ern‐central Ecuador where Quaternary arc vol‐
canism is restricted to the area north of c. 2.5°S, and Late Miocene arc volcanic formations cover small areas between c. 2.5° and 4°S (Gutscher et al. 1999; Chapter 2). The Andean chain hosts the bulk of Tertiary arc magmatic products and splits into a western and eastern Cordillera in central‐northern Ecuador, which are separated by a number of elongated basins referred to as Interandean Depression (Litherland et al. 1994; Winkler et al. 2005). In southern Ecuador, the Andean structural NNE trend is disrupted where the Western Cordillera swings towards the Gulf of Guayaquil and is re‐
placed by the El Oro micro‐continental block which underwent clockwise rotation during the Cretaceous‐Tertiary resulting in an arc‐transverse structural trend at the present day (Mitouard et al. 1990; Litherland et al. 1994). Late Oligocene‐Miocene arc volcanics are mostly eroded in the Western Cordillera of central‐
northern Ecuador such that their deeper‐seated plutonic equivalents are unroofed (Chapter 2). These plutons are aligned along major fault zones of several 100’s km strike length which extend down to mid‐ to deep crustal levels where they are defined by 35°E dipping fault planes (Guillier et al. 2001). The possible eastward continuation of Late Oligocene‐Miocene arc magmatism is concealed below Quaternary arc volcanic cover sequences of the Interandean Depression (Fig. 1; Chapter 2). Miocene ore deposits of the Western Cordillera and its western foothills mostly repre‐
sent moderately to deeply eroded porphyry sys‐
tems such as Junin, Balsapamba, and Telimbela, partly associated with minor epithermal minerali‐
zation (Prodeminca 2000a). Older porphyry sys‐
tems have not been described, but a number of Au‐rich Eocene volcanic‐hosted massive sulfide deposits occur (Chiaradia & Fontboté 2001; Chiaradia et al. 2008). The Late Oligocene to Early Miocene Saraguro Group constitutes the major outcrop unit of the southern Ecuadorian Sierra north of the Piñas‐
Portovelo fault (Fig. 2). It overlaps in age with the Calipuy Group in northern Peru and is partly overlain by volcaniclastic‐sedimentary formations of the Cuenca and associated intramontane ba‐
sins, and by Mid‐ to Late Miocene arc volcanic formations (Sta. Isabel, Quimsacocha, Tarqui; Chapter 2). As in northern Peru, extensive Qua‐
ternary volcanic cover sequences are absent in southern Ecuador creating a favorable erosion level for the exposure of Miocene mineralization (Fig. 2). The Saraguro Group volcanics are punctured by numerous intrusions including the major Cangre‐
jos‐Zaruma intrusive belt, and host a large num‐
ber of epithermal and porphyry Cu deposits which are, as shown below, mainly of Miocene age (Fig. 2). The ore deposits of southern Ecuador define two main districts referred to as Azuay and El Oro districts, respectively (Prodeminca 2000a, b; Fig. 1, 2). Additional pre‐Miocene Terti‐
ary mineralization in the southern Ecuadorian Sierra was not identified in the present study but cannot be ruled out. Pre‐Tertiary mineralization is evidenced by the Late Cretaceous Curiplaya por‐
phyry intrusions in SW Ecuador (Fig. 1; Chapter 2), and the highly prolific Jurassic period of min‐
eralization in the Eastern Cordillera, including, amongst others, the Fruta del Norte, Mirador, and Nambija deposits (Gendall et al. 2000; Stew‐
art & Leary 2007; Chiaradia et al. 2009b). Local geology of Miocene Ecua‐
dorian ore deposits investigated in this study Seven ore deposits of northern, central and, mainly, southern Ecuador were sampled for Re‐
Os molybdenite dating in this study. These 69
70 comprise deposits with porphyry‐style minerali‐
zation (from north to south: Junin, Telimbela, Balsapamba, Chaucha, Gaby, Cangrejos) and one breccia‐related epi‐ to mesothermal deposit (Tres Chorreras) whose metallogenic classification is not entirely clear (Prodeminca 2000a). In addi‐
tion, hydrothermal titanite was sampled for U‐Pb dating at the Gaby porphyry system. General geo‐
logical features of these deposits are summarized in Table 1. Typical alteration and mineralization characteristics of these deposits are shown in Figure 3, and geological maps are provided for a number of key deposits where more detailed geochronologic studies were carried out (Fig. 4‐
7). The Junin Cu‐Mo porphyry system is hosted by the Mid‐Miocene Apuela batholith (Fig. 1; Chap‐
ter 2). It occurs in the center of a belt of three porphyry deposits (including El Pacto to the SW and Cuellaje to the NE; plus a meso‐ (?) to epi‐
thermal Au deposit at El Corazon) which are aligned in NE directions parallel to the Chimbo‐
Toachi shear zone, and are collectively referred to as Imbaoeste district (MMAJ/JICA 1998; Prodeminca 2000a; Micon 2005a; Chapter 2). The Junin prospect comprises a well developed zone of phyllic‐potassic alteration partly extending downwards to 600 m depth, centered on multiple hornblende granodiorite porphyry dikes of vari‐
able thickness striking NNE to ENE and dipping 45‐70° to the SE (Fig. 4; Salazar 2007). Local struc‐
tures show a major ~NE trend and secondary N‐
NW structures which were repeatedly active at pre‐, syn‐, and postmineral times, and are in‐
ferred to have controlled porphyry dike em‐
placement by facilitating local dilation (Micon 2005a). The Balsapamba and Telimbela Cu‐Mo porphyry systems define the Bolivar district in central Ec‐
uador and occur in the western foothills of the Western Cordillera (Fig. 1); they represent deeply eroded porphyry systems hosted by various fa‐
cies of the central Ecuadorian Oligocene‐Miocene batholith (MMAJ/JICA 1991; Prodeminca 2000a; Chapter 2). Hydrothermal alteration (mainly potassic ± sodic‐calcic) is centered on multiple hornblende quartz‐diorite porphyry dikes which are aligned with mainly NE‐, but also N‐, NW‐, and ENE‐trending structures (Prodeminca 2000a). Local advanced argillic alteration has been ob‐
served in places (MMAJ/JICA 1991). Overall, Cu mineralization related to the Balsapamba por‐
phyry systems (mainly the El Torneado zone) seems to have been mostly eroded, whereas various exploration targets in the Telimbela por‐
phyry system have a higher mineralization poten‐
tial, especially in brecciated areas (MMAJ/JICA 1991). The Chaucha Cu‐Mo porphyry system is one of the earliest described porphyry Cu deposits in Ecuador (e.g., Goossens & Hollister 1973). It is situated next to the major Bulubulu fault system at the SE end of the Mid‐Miocene Chaucha ba‐
tholith and comprises at least two major por‐
phyry intrusions (Tunas and Gur‐Gur) hosted by pre‐Tertiary metapelites, Saraguro Group volcan‐
ics, and older intrusive phases of the Chaucha batholith (Figs. 2, 6; Prodeminca 2000a; Micon 2005b; Chapter 2). The spatial distribution of hy‐
drothermal alteration reflects the trends of prin‐
cipal tectonic structures and overall affects an area of several km2; highest ore grades are en‐
countered in zones of transitional potassic‐phyllic alteration both in porphyry stocks and batholith host units (Micon 2005b). The Gaby‐Papa Grande Au‐Cu porphyry system comprises multiple Early Miocene porphyry and phaneritic intrusions (stocks and dikes) emplaced in oceanic plateau basalts (Pallatanga Unit), oc‐
curring at a short distance to the epithermal Bella Rica Au vein system (Fig. 2, 7; Prodeminca 2000a; Srivastava et al. 2008; Chapter 2). Gold (‐Cu) por‐
phyry mineralization seems to be associated with sodic‐calcic alteration (mostly as free Au) and is particularly well developed in previously frac‐
tured porphyry intrusions and hydrothermal breccias which show an overall NW distribution trend (Srivastava et al. 2008). Local high‐grade mineralization is structurally controlled and asso‐
ciated with phyllic alteration (e.g., Tama vein; Fig. 7; Srivastava et al. 2008). The Gaby and Papa Grande porphyry systems are separated by the E‐
W striking Guanache normal fault resulting in a deeper exposure level of the Gaby relative to the Papa Grande sector (Prodeminca 2000a). Tres Chorreras constitutes the northernmost end‐
member of a series of tourmaline‐bearing breccia pipe‐related deposits emplaced along the NE‐
trending La Tigrera fault in the southern 71
Figure 3: Typical alteration and mineralization features of Late Tertiary porphyry systems of Ecuador (A‐D = macro‐; E‐H = micro‐photographs). A – Multiple veinlets of cp‐qtz and qtz‐ms‐py‐mo cross‐cutting granodiorite porphyry with per‐
vasive potassic and phyllic alteration (Junin). B – Hydrothermal breccia with subangular hbl‐plag porphyry clasts (with sodic‐calcic alteration) and bt‐qtz cement (related to Tama vein; Gaby). C – Multiple mt and qtz‐mt veinlets with ep haloes cross‐cutting hbl qtz‐diorite porphyry (Cangrejos). D – Reopened mt‐qtz vein filled with later qtz and ser halo cross‐cutting tonalite with bt‐chl alteration, plus multiple thin cp‐qtz veinlets (Chaucha batholith at Tunas). E – hbl‐
bearing granodiorite porphyry with pervasive potassic alteration where bt flakes completely replace hbl phenocrysts
72 Ecuadorian Sierra. Its position corresponds to the projected intersection of three regional linea‐
ments, namely the La Tigrera fault, the SE‐
trending Galena fault, and the NNE‐trending Bu‐
lubulu fault (Fig. 2; Prodeminca 2000a). Lithologi‐
cal units at Tres Chorreras comprise a number of diorite‐granodiorite intrusions emplaced in silicic Saraguro Group volcanics and volcaniclastics, which are associated with several breccia pipes and a subcircular to irregularly‐shaped agglomer‐
ate‐filled structure interpreted as a large dia‐
treme (Prodeminca 2000a). Gold mineralization in quartz veinlets is partly hosted by various breccia bodies and by the diatreme structure. In addition, a younger set of polymetallic veins oc‐
cur (Prodeminca 2000a). While broadly classified as meso‐ to epithermal mineralization (Prodeminca 2000a), some of the deposit's geo‐
logic features such as tourmaline‐bearing hydro‐
thermal breccias are typically porphyry‐related (e.g., Seedorff et al. 2005), whereas other fea‐
tures such as polymetallic vein mineralogy are similar to Cordilleran vein‐type deposits (e.g., Fontboté & Bendezú 2009). The Late Oligocene‐Early Miocene Cangrejos Au‐
Cu porphyry system occurs at the western end of the Cangrejos‐Zaruma intrusive belt (Fig. 2; Chap‐
ter 2). It comprises multiple nested intrusions punctured and intruded by a number of porphyry dikes and breccia pipes (Potter 2004). Gold is as‐
sociated with sulfides or occurs in quartz veinlets whose distribution is structurally controlled and includes all intrusive lithologies; highest Au grades are associated with quartz‐tourmaline veinlets (Potter 2004). Sampling and analytical tech‐
niques Sampling details are listed in Table 2 and further illustrated on Figures 4 to 7. Molybdenite sam‐
ples used for Re‐Os datation are shown in Figure 8. Pure molybdenite concentrates of 10‐60 mg/sample were obtained from massive molyb‐
denite or quartz‐molybdenite veinlets of samples listed in Table 2 using a microdrill, followed by handpicking to purify the concentrates. Where molybdenite occurred as fine‐grained flakes in the matrix of hydrothermal breccias (Gaby, Balsapamba; Fig. 9) samples were crushed to <300 μm and washed to remove clay particles, followed by molybdenite handpicking from the heavy (> 3.32 g/cm3) non‐magnetic mineral frac‐
tion > 80 μm. Rhenium and Os were separated at the University of Arizona according to the proce‐
dures described in Barra et al. (2003, 2005). Weighted molybdenite fractions were spiked with 185Re and 190Os and dissolved in a Carius tube using 8 ml inverse aqua regia (3 ∙ 16N HNO3 + 1 ∙ 10N HCl); 2‐3 ml of hydrogen peroxide (30%) were added to the mixture to ensure complete sample oxidation and spike equilibration. The tube was heated to 240°C for c. 8 h, and the solu‐
tion subsequently treated in a two‐stage distilla‐
tion process for Os separation (Nägler & Frei 1997). Osmium was further purified using a mi‐
crodistillation technique, similar to that of Birck et al. (1997), and loaded on Pt filaments with Ba(OH)2 to enhance ionization. After Os separa‐
tion, the remaining acid solution was dried and later dissolved in 0.1 N HNO3. Rhenium was ex‐
tracted and purified through a two‐stage column using AG1‐X8 (100–200 mesh) resin and loaded on Pt filaments with Ba(SO)4. Samples were analyzed by negative thermal ioni‐
zation mass spectrometry (Creaser et al. 1991) on a VG 54 mass spectrometer at the University of Arizona. Rhenium and Os were measured with Faraday collectors. Molybdenite ages were calcu‐
lated using an 187Re decay constant of 1.666 ∙10‐11 year‐1 (Smoliar et al. 1996). Errors are reported at the 2σ level and cmprise the propa‐
gated uncertainties of the Re decay con‐
Figure 3 (caption continued from previous page): and form "shreddy" disseminations in the porphyry matrix (Gur‐Gur porphyry, Chaucha). F – Sodic‐calcic alteration‐related vein with ttn, act, ep, and po‐cp cross‐cutting plag‐hbl porphyry (Gaby). G – Plag‐qtz porphyry with chl background alteration, cross‐cut by ep‐py‐cp veinlet (Telimbela). H – Porphyritic granodiorite with pervasive silicification and ms alteration (Gur‐Gur porphyry, Chaucha. Mineral abbreviations: ab –
albite, act – actinolite, bt – biotite, chl – chlorite, cp – chalcopyrite, ep – epidote, hbl – hornblende, jsp – jasper‐like silica; mo – molybdenite, mt – magnetite, ms – muscovite, plag – plagioclase, po – pyrrhotite, py – pyrite, qtz – quartz, rt – rutile, ser – sericite, sl ‐ sphalerite; tm – tourmaline, ttn – titanite. ). Scale bar is 2 cm for macro‐, and 1 mm for mi‐
cro‐photographs. 73
stant (0.31%), spike calibration for 185Re (0.08%) and 190Os (0.15%), and individual weighting and analytical random errors. Weighted mean ages were calculated using the Isoplot v3.31 Excel macro (Ludwig 2003). Hydrothermal titanite forms part of the sodic‐
calcic alteration assemblage at the Gaby por‐
phyry system (Tab. 2). Titanite from the crushed and milled <400 μm grain size fraction of sample E05077 was separated using standard Wilfley ta‐
ble and heavy mineral (> 3.32 g/cm3) separation techniques, and handpicked from the slightly magnetic fraction (0.8‐1.25 A @ 20° side tilt) us‐
ing a Frantz magnetic separation table. The ti‐
tanite fraction underwent a bulk two‐step wash‐
ing process at 140°C (30 m in. each) using (1) a mixture of concentrated HF and 7N HNO3, and (2) 6N HCl, followed by rinsing in ultrapure H2O and acetone. Titanite multi‐grain fractions (n = 4‐7) were spiked using a mixed 205Pb‐233U‐235U spike solution, and were dissolved in 63 μl concen‐
trated HF with a trace of 7N HNO3 at 110°C for seven days. Uranium and Pb were separated us‐
ing an HCl‐based anion exchange chromatogra‐
phy, and loaded individually on separate Re fila‐
ments using the Si‐gel technique (Gerstenberger & Haase 1997). Measurement routines on a Tri‐
ton thermal ionization mass spectrometer at the University of Geneva were identical to those out‐
lined for zircon analysis in Chapter 2. Total analytical common Pb (Pbc) was attributed to both laboratory blank (isotopically constrained by repeated measurements as 206Pb/204Pb = 17.87±0.36, 207Pb/204Pb = 15.16±0.34, 208Pb/204Pb = 36.75±1.11) and Pbc included in titanite (iso‐
topic composition estimated at t = 20 Ma accord‐
ing to Stacey & Kramers 1975). Lab blanks were highly variable (15‐81 pg), but constant in iso‐
topic composition; therefore, individual propor‐
tions of titanite Pbc vs. Pbc introduced by labora‐
tory contamination were calculated assuming a constant Pbc concentration in titanite which was obtained iteratively by balancing the weight of the titanite fraction vs. the total amount of Pbc, yielding an average titanite Pbc concentration of 229±14 pg/mg. The uncertainties of spike and blank Pb isotopic composition, mass fractionation correction, and tracer calibration were propa‐
74 gated to the final uncertainties of isotopic ratios and ages of each individual analysis. Uncertain‐
ties in the decay constants of 238U and 235U (Jaffey et al. 1971) were propagated separately and added quadratically to the weighted mean age. Concordia plots and weighted mean age calcula‐
tions were prepared using the Isoplot v.3.31 Excel macro of Ludwig (2003). All uncertainties and error ellipses are reported as 2σ and weighted mean 206Pb/238U ages are presented at 95% con‐
fidence level. Results Table 3 shows Re‐Os data for the ten molyb‐
denite concentrates analyzed in this study. Total Re and 187Os concentrations range between 35‐
1019 ppm and 4.7‐250 ppb, respectively. Two Re‐ Os molybdenite ages of 6.63±0.04 Ma and 6.13±0.03 Ma were obtained for quartz‐
molybdenite veinlets at the Junin porphyry sys‐
tem. This veinlet type is related to potassic or transitional potassic‐phyllic alteration at Junin (Salazar 2007). The new Re‐Os molybdenite age at Telimbela (19.2±0.1 Ma) was obtained on a molybdenite‐quartz veinlet related to potassic alteration. A Re‐Os molybdenite age of 21.5±0.1 for the El Torneado zone of the Balsapamba plu‐
ton relates to molybdenite as part of a hydro‐
thermal breccia matrix consisting of mineral phases of a potassic alteration assemblage (Tab. 2). Two Re‐Os molybdenite ages at Chaucha comprise the Tunas‐Naranjos sector (9.92±0.05 Ma) and the Gur‐Gur sector (9.5±0.2 Ma; Fig. 6). The former age was obtained on molybdenite associated with potassic‐phyllic alteration hosted by biotite‐bearing granodiorite, and the latter age was obtained on molybdenite hosted by grano‐
diorite porphyry with phyllic alteration. A Re‐Os molybdenite age of 20.6±0.1 Ma was obtained at the Gaby porphyry system where molybdenite forms part of a hydrothermal breccia matrix (Tab. 2). At Tres Chorreras we dated molybdenite as part of a hydrothermal breccia matrix (12.93±0.07 Ma), and as part of a massive poly‐
metallic vein (12.75±0.07 Ma). Finally, a molyb‐
denite‐quartz veinlet associated with sodic‐calcic alteration at Cangrejos gave a Re‐Os molybdenite age of 23.5±0.1 Ma. Table 2: Description of samples used for molybdenite Re-Os and titanite U-Pb datation
Deposit/sample Location/drill core
Description
Junin
E06194
E06199
Telimbela
E07037
Balsapamba
E08003
Chaucha
E07006 (Naranjos
sector, adjacent to
Tunas porphyry)
E06175 (Gur-Gur
porphyry)
Gaby-Papa
Grande
E05075
E05077
Tres Chorreras
E07010
E07012
Cangrejos
E06065
35050 N, 761383 E;
0° 19' 1'' N, 78° 39' 6'' W;
MJJ-29 @ 305m
hbl granodiorite porphyry (potassic, overprinted by phyllic alteration);
qtz-mo (-cp) veinlet with fine-grained mo flakes
35050 N, 761383 E;
0° 19' 1'' N, 78° 39' 6'' W;
MJJ-29 @ 498m
hbl granodiorite porphyry (potassic, overprinted by phyllic alteration);
mo-qtz veinlets with coarse-grained mo flakes
9817260 N, 705670 E;
1° 39' 9'' S, 79° 9' 5'' W;
MJE-9 @ 49m
hbl tonalite (potassic, overprinted by propylitic alteration); multiple
mo-qtz veinlets with coarse-grained mo flakes
9808050 N, 707840 E;
1° 44' 9'' S, 79° 7' 54'' W;
MJE-3 @ 42m
brecciated hbl granodiorite (potassic-phylic alteration) with breccia
matrix (+veinlets?) of chl-bt-qtz-mt with mo-cp-py; fine-grained mo
flakes in breccia matrix
9676800 N, 676140 E;
2° 55' 23'' S, 79° 24' 55''
W; NA-30 @ 53m
altered host tonalite (potassic-phyllic) with intense qtz, qtz-cp, cp-mtqtz veining, and thick qtz vein with mo concentrated at vein margins
(fine-grained mo flakes)
9676700 N, 677800 E;
2° 55' 26'' S, 79° 24' 1'' W;
core-5 @ 80m
altered granodiorite porphyry (phyllic) with py-qtz and qtz-py-mo
veinlets (fine-grained mo flakes)
9661850 N, 643400 E;
3° 3' 31'' S, 79° 42' 35'' W;
GD-08 @ 151m
hydrothermal breccia with altered (qtz-ser), subangular porphyry
clasts; breccia matrix comprises qtz, po, cp±mo, and goe, hm, jar
(later oxidation?); mo as fine-grained flakes
9661850 N, 643400 E;
3° 3' 31'' S, 79° 42' 35'' W;
GD-08 @ 340m
hbl-plag porphyry with strong pervasive Na-Ca alteration including
act, chl, ep, ttn, and sulfides (mainly po); anhedral-euhedral ttn
grains of c. 50-500 μm size in porphyry matrix or replacing hbl (along
with other minerals)
9650150 N, 663591 E;
3° 9' 51'' S, 79° 31' 40'' W
hydrothermal breccia with subangular clasts of altered volcanics and
matrix of tm, mt (repl. by hm + goe), mo, jsp; mo as coarse-grained
flakes
massive polymetallic vein with cp-mo-jsp-sl; host rock = completely
replaced by clay minerals (pyrophyllite?); mo as coarse-grained
flakes
9650052 N, 663543 E;
3° 9' 54'' S, 79° 31' 42'' W
9614000 N, 633200 E;
3° 29' 29'' S, 79° 48' 3'' W
bt-bearing qtz-diorite with weak pervasive Na-Ca alteration; abundant qtz veinlets and single mo-qtz veinlet (coarse-grained mo
flakes)
Same mineral abbreviations as in Table 1, plus goe (goethite), jar (jarosite). Coordinates as PSAD-56 projection.
Hydrothermal titanite is associated with sodic‐
calcic alteration at the Gaby porphyry system where a U‐Pb titanite age of 20.17±0.16 Ma was obtained (Fig. 9; Tab. 4). High scatter between individual 207Pb/235U ages for Gaby titanite sug‐
gests an imperfect characterization of the Pbc isotopic composition, but the homogeneous dis‐
tribution of 206Pb/238U ages implies that 206Pb/238U age systematics were not significantly affected by this issue. 75
Discussion Integration of Re‐Os molybdenite and U‐Pb titanite ages into the geochro‐
nologic framework of individual ore deposits The new Re‐Os molybdenite ages at Junin (6.63±0.04 and 6.13±0.03 Ma) are significantly younger than the U‐Pb zircon age of 9.01±0.06 Ma of a hornblende granodiorite porphyry dike from the same drill core (Fig. 5; Chapter 2). As the relative Re‐Os age difference of 0.5 m.y. is out‐
side the maximum life span of a moderately‐
sized, single intrusion‐driven hydrothermal sys‐
tem at shallow depth (e.g., Marsh et al. 1997) this age distribution suggests that widespread potas‐
sic‐phyllic alteration is related to several (at least two) post‐9 Ma porphyry systems. Published whole rock and biotite/hornblende K‐Ar ages of Junin porphyry intrusions are 7.9±0.3 Ma, 7.3±0.3 Ma, 6.1±0.2 Ma, and 5.9±0.1 Ma (MMAJ/JICA 1992; Prodeminca 2000a). The youngest Re‐Os age thus overlaps (within error) with and con‐
firms the youngest K‐Ar ages, whereas the 6.6 Ma Re‐Os molybdenite age evidences an additional hydrothermal pulse previously not detected by K‐
Ar dating. In agreement with geological and petrographic studies (Salazar 2007), the variable age range reflects multiple intrusive events and hydrothermal systems at Junin. Following a major period of host batholith construction from c. 19 to 12 Ma (Chapter 2, and references therein) re‐
peated porphyry dike emplacement associated with several hydrothermal systems occurred be‐
tween 9 and 6 Ma. Previously obtained K‐Ar ages of 7.9 and 7.3 Ma (overlapping with each other within error) might either reflect an additional intrusive event at that time, or might relate to older intrusive events (such as porphyry dike em‐
placement at 9.01 Ma) and subsequent distur‐
bance of the K‐Ar isotopic system by younger in‐
trusive/hydrothermal pulses. Figure 4: Geological map of the Junin Cu‐Mo±Ag porphyry system (adapted from MMAJ/JICA 1998).
76 Table 3: Re-Os data for molybdenite of Miocene Ecuadorian ore deposits
Deposit/sample
Junin
E06199
E06194
Telimbela
E07037
Balsapamba
E08003
Chaucha
E07006
E06175
Gaby-Papa Grande
E05075
Tres Chorreras
E07010
E07012
Cangrejos
E06065
weight
[mg]
Total Re
[ppm]
187
Re
[ppm]
187
Os
[ppb]
39
16
294.4
408.8
184.3
255.9
18.8
28.3
6.13 ± 0.03
6.63 ± 0.04
60
312.8
195.8
62.6
19.2 ± 0.1
14
580.1
363.1
130.4
21.5 ± 0.1
30
11
354.6
70.4
222.0
44.1
36.7
7.0
9.92 ± 0.05
9.5 ± 0.2
44
442.9
277.3
95.0
20.6 ± 0.1
55
52
641.0
35.3
401.3
22.1
86.4
4.7
12.93 ± 0.07
12.75 ± 0.07
50
1019
637.9
249.9
Propagated total age uncertainties (c. 0.5%) include uncertainties in the Re decay constant (0.31%),
190
age ± 2б
[Ma]
23.5 ± 0.1
185
Re (0.08%) and
Os (0.15%) spike calibration, weighting, and analytical random errors.
Weighted mean ages were calculated using the Isoplot v.3.31 Excel macro (Ludwig 2003).
A Re‐Os molybdenite age of 19.2±0.1 Ma over‐
laps with several K‐Ar ages obtained on various facies of the Telimbela pluton (19.1‐19.4 Ma; MMAJ/JICA 1989; McCourt et al. 1997) where the total age range of the Telimbela pluton is 25.5‐
14.5 Ma (K‐Ar hornblende, biotite, and whole rock data; MMAJ/JICA 1989, 1991; McCourt et al. 1997; Chapter 2). This suggests that formation of the porphyry system occurred after some 6 m.y. of pluton construction, and was still followed by younger plutonic activity. A K‐Ar age of 15.7±1.0 Ma was obtained on a “quartz‐porphyry” (MMAJ/JICA 1991) possibly indicating further de‐
velopment of porphyry systems towards the end of pluton assembly. Our new Re‐Os molybdenite age of 21.5±0.1 Ma from the El Torneado zone of the northern Balsa‐
pamba pluton significantly predates a previously obtained Re‐Os molybdenite age of 19.7±0.3 Ma from the same area (Chiaradia et al. 2004). It is identical with the U‐Pb zircon age of the major pluton lithology (hornblende‐bearing granodio‐
rite; 21.5±0.1 Ma; Chapter 2), and overlaps within error with the U‐Pb zircon age of a quartz‐diorite porphyry dike (21.2±0.2 Ma; Chapter 2); both reference ages were also obtained within the El Torneado zone of the northern Balsapamba plu‐
ton. These ages are significantly younger than previously published K‐Ar ages of the Balsapamba pluton (33.1‐25.7 Ma; MMAJ/JICA 1989; McCourt et al. 1997) and show that the Balsapamba and Telimbela plutons, along with the spatially asso‐
ciated plutons of Chaso Juan, Las Guardias, El Co‐
razon, and La Industria, were formed at broadly similar times and constitute an intrusive complex of batholith dimension (cf. Chapter 2). The inte‐
grated geochronologic results for the Balsapamba and Telimbela plutons demonstrate that multiple porphyry‐related hydrothermal pulses occurred in a relatively short time span during the Early Miocene, following a multi‐m.y. history of batho‐
lith construction. Molybdenite associated with hydrothermal al‐
teration at the Tunas‐Naranjos and Gur‐Gur por‐
phyry systems at Chaucha yields different Re‐Os ages (9.92±0.05 Ma vs. 9.5±0.2 Ma). Both por‐
phyry systems comprise different intrusive lithologies, alteration characteristics, and host rocks, and occur at >2 km distance to each other (e.g., Prodeminca 2000a; Fig. 6). Combined with 77
the Re‐Os relative age difference of 0.4 m.y. (again, as in Junin, well outside the maximum life span of a single intrusion‐driven hydrothermal system at shallow depth) these characteristics suggest that the Tunas and Gur‐Gur intrusions define two distinct porphyry systems separated in time and space, although they might ultimately be related to the same parental magmatic system at depth. The 9.92±0.05 Ma Re‐Os age obtained on molyb‐
denite hosted by biotite‐bearing granodiorite with potassic‐phyllic alteration dated at 14.8±0.1 Ma (U‐Pb zircon; Chapter 2) closely approaches but pre‐dates the U‐Pb zircon age (9.79±0.03 Ma; Chapter 2) of a close‐by granodiorite porphyry dike in the Tunas sector at Chaucha whose em‐
placement might have been associated with a hydrothermal porphyry system. The age differ‐
ence between the non‐overlapping zircon and molybdenite ages might be due to an overestima‐
tion of the molybdenite age by minor alteration‐
induced Re‐loss (e.g., Barra et al. 2003), underes‐
timated analytical uncertainties, or, possibly, an underestimation of the zircon age by subtle ra‐
diogenic Pb loss (Chapter 2). Similar to Junin and Balsapamba‐Telimbela, porphyry deposit forma‐
tion at Chaucha postdates the major batholith construction period. Multiple intrusive phases are present at Gaby and Papa Grande (Fig. 7) where the main lithologies, represented by hornblende‐plagioclase porphyry stocks, have been dated at 20.26±0.07 Ma (Gaby) and 19.89±0.07 Ma (Papa Grande; U‐Pb zircon; Chapter 2). At Gaby, the dated main porphyry body is additionally cut by multiple porphyry dikes. The age of hydrothermal titanite (20.17±0.16 Ma) associated with sodic‐calcic al‐
teration at Gaby overlaps within error with the age of the main porphyry intrusion, consistent with a close relationship between intrusion em‐
placement and fluid circulation causing sodic‐
calcic alteration. Figure 5: Simplified lithology‐alteration drill core log of Junin core MJJ‐29 showing multiple porphyry intrusive phases and locations of samples dated in this study
and in Chapter 2. The mining company's internal lithologic classification, as applied here, does not strictly correspond to the classification used throughout this study such that minor differences in the distribution of intrusive bodies exist. For example, "coarse‐grained quartz‐feldspar porphyry" corresponds to "hornblende granodiorite porphyry" in our
(BGS‐based) classification scheme. Hornblende granodiorite porphyry was dated at 9.01 Ma (E07032; Chapter 2) and seems to represent the oldest porphyry intrusion at Junin dated so far, where the wall rock (Apuela batholith) is
>12.9 Ma (Chapter 2). Molybdenite Re‐Os and various K‐Ar ages indicate younger porphyry intrusions were emplaced until c. 6 Ma. Adapted from Salazar (2007). 78 In contrast, our new Re‐Os molybdenite age (20.6±0.1 Ma) is older than the previously dated porphyry intrusion. The age difference might be due to an overestimate of the Re‐Os molybdenite age caused by Re loss of molybdenite which, for example, might accompany crystallographic transformations from 3R to 2H polytypes (McCandless et al. 1993; see also Barra et al. 2003). Alternatively, one or more additional pre‐
20.6 Ma porphyry intrusions might be present at Gaby, which exsolved fluids responsible for mo‐
lybdenite precipitation. Earlier intrusive activity at Gaby is evidenced by the occurrence of an un‐
dated tonalite intrusion closely associated with and (according to field relationships) predating the porphyry intrusions (Fig. 7). The two Re‐Os molybdenite ages at Tres Chor‐
reras (12.93±0.07 Ma and 12.75±0.07 Ma) do not overlap within error. The older age dates the tim‐
ing of hydrothermal brecciation possibly related to the emplacement of spatially associated por‐
phyry intrusions for which no ages are available at present (Prodeminca 2000a). The younger age dates the timing of polymetallic mineralization. Both ages are significantly younger than their volcanic host units (30.7±0.7 Ma; Chapter 2). The small age difference (180 k.y.) between porphyry‐
style and polymetallic mineralization might indi‐
cate that both events ultimately relate to the same magmatic‐hydrothermal system, but more detailed geologic studies of the Tres Chorreras deposit are required before any qualified conclu‐
sions can be drawn. A quartz‐diorite of the Cangrejos intrusive com‐
plex has a U‐Pb zircon age of 26.0±0.7 Ma (Chap‐
ter 2) and is intruded by plagioclase‐hornblende porphyry (Potter 2004). The new molybdenite age of 23.5±0.1 Ma might thus be related to a hydrothermal system generated by the post‐26 Ma porphyry intrusion. These ages are signifi‐
cantly older than a K‐Ar age of 16.9±0.2 Ma ob‐
tained on the Paccha intrusion in the center of the Cangrejos‐Zaruma intrusive belt (Fig. 2; Pratt et al. 1997) which had previously been proposed as an age reference for the Cangrejos porphyry system (Potter 2004). In contrast, these ages are broadly similar to U‐Pb zircon ages of intrusions in the Portovelo‐Zaruma mining district (20.7±0.9 Ma and 24.0±0.1 Ma; Chapter 2) at the western end of the Cangrejos‐Zaruma intrusive belt. 79
Figure 6: Geological map of the Chaucha Cu‐Mo porphyry system showing locations of dated samples. Samples for U‐
Pb zircon datation (Chapter 1) were collected from surface outcrop exposure whereas samples for Re‐Os molybdenite datation (this study) are drill core samples (cf. Tab. 2). Adapted from Micon (2005b). A quartz‐diorite of the Cangrejos intrusive com‐
plex has a U‐Pb zircon age of 26.0±0.7 Ma (Chap‐
ter 2) and is intruded by plagioclase‐hornblende porphyry (Potter 2004). The new molybdenite age of 23.5±0.1 Ma might thus be related to a hydrothermal system generated by the post‐26 Ma porphyry intrusion. These ages are signifi‐
cantly older than a K‐Ar age of 16.9±0.2 Ma ob‐
tained on the Paccha intrusion in the center of the Cangrejos‐Zaruma intrusive belt (Fig. 2; Pratt et al. 1997) which had previously been proposed as an age reference for the Cangrejos porphyry system (Potter 2004). In contrast, these ages are broadly similar to U‐Pb zircon ages of intrusions 80 in the Portovelo‐Zaruma mining district (20.7±0.9 Ma and 24.0±0.1 Ma; Chapter 2) at the western end of the Cangrejos‐Zaruma intrusive belt. Magmatic characteristics of the Mio‐
cene metallogenic belt of Ecuador Miocene Ecuadorian ore deposits investigated in this study are always intimately associated with intrusive activity (e.g., Fig. 2; Chapter 2). Several porphyry Cu deposits in Ecuador (Junin, Balsa‐
pamba‐Telimbela, Chaucha) are associated with the final pulses of batholith‐scale intrusive sys‐
tems, which record protracted periods of precur‐
sor magmatism over several million years. To visualize this, we have plotted the distribution of intrusive and mineralization‐/alteration‐related radiometric ages for several Ecuadorian arc seg‐
ments (Fig. 10). All major batholith systems of Ecuador associated with ore deposits show a similar pattern where porphyry deposits form c. 5 m.y. (Chaucha), 10‐13 m.y. (Junin), or 13‐15 m.y. (Balsapamba‐Telimbela) after initialization of ba‐
tholith magmatism. A similar precursor intrusive history might be inferred for the Cangrejos por‐
phyry system, although available geochronologic data is too scant to quantify this. Similar observa‐
tions have been made elsewhere in the Miocene metallogenic belts of Chile and Peru (e.g., Sillitoe 1988) and in the Jurassic metallogenic belt of Ec‐
uador (Chiaradia et al. 2009b). The only potential deviation from this pattern is represented by the Gaby porphyry system, where geochronologic evidence points to a rather short‐lived intrusive system and directly associated large intrusive bodies are absent. Multi‐million year batholith assembly signals effi‐
cient channeling of arc magmas ascending through the crust resulting in large, repeatedly replenished mid‐ to shallow crustal magmatic systems; catastrophic, caldera‐forming ignimbrite eruptions often accompany voluminous batho‐
lith‐related magmatism (e.g., Bachmann et al. 2007). Thus, shallow crustal batholith sites repre‐
sent a potentially favorable environment for the formation of porphyry‐related ore deposits as they provide large volumes of magma and ther‐
mal energy to drive hydrothermal systems (e.g., Cline & Bodnar 1991). However, intense shallow crustal magmatism might have negative implica‐
tions for mineralization (e.g., due to dispersed or catastrophic volatile loss instead of focused fluid exsolution) or its preservation (e.g., destruction of mineralization by subsequent intrusive pulses). During the waning stages of batholith assembly, on the other hand, less vigorous magma replen‐
ishment and downwards migration of the focus of magmatic activity might represent a favorable tectonomagmatic environment to form and pre‐
serve porphyry‐related ore deposits where pro‐
gressive melt volatile‐enrichment at mid‐crustal levels takes place (e.g., Rohrlach & Loucks 2005; Chiaradia et al. 2009a). In this context, the ton‐
nage of potentially formed ore deposits is not expected to directly correlate with the total dura‐
tion of precursor intrusive magmatism (e.g., Har‐
ris et al. 2004; Barra et al. 2005). Systematic across‐arc variations with respect to the timing of magmatism and mineralization do not seem to exist in southern Ecuador (Fig. 2, 10). Instead, both magmatism and metallogenesis seem to span the whole width of the arc segment at a given time, in agreement with an inferred period of arc broadening in the Early Miocene (Chapter 2). Mid‐Miocene intrusion emplacement (c. 16 Ma; Chapter 2) and advanced argillic altera‐
tion (15.4 Ma; K‐Ar alunite; Prodeminca 2000b) at the El Mozo high‐sulfidation epithermal deposit, situated at the easternmost margin of the Mio‐
cene metallogenic belt, broadly coincide in time with Mid‐Miocene Chaucha batholith magmatism (15‐10 Ma; Chapter 2), polymetallic mineraliza‐
tion at Tres Chorreras (12.9‐12.8 Ma), and mag‐
matism (15.7 Ma, U‐Pb zircon) and hydrothermal alteration (18.9 Ma, Ar‐Ar sericite) at the Rio Blanco low‐sulfidation deposit (Bineli Betsi 2007), all situated at the western side of the belt. The timing of mineralization may differ profoundly at a given position within the metallogenic belt: for example, the neighboring (at c. 40 km distance) Gaby and Chaucha porphyry systems formed at c. 20 Ma and 10 Ma, respectively. These considera‐
tions corroborate recent results of Noble et al. (2004) for the northern‐central Peruvian metal‐
logenic belt where mineralization was partly coe‐
val at the western and eastern belt extremities. Structural characteristics of the Mio‐
cene metallogenic belt of Ecuador The principal spatial distribution of Tertiary intru‐
sions in Ecuador mimics the major upper plate structures (Chapter 2); these include the Chimbo‐
Toachi shear zone in central‐northern Ecuador (associated with the Apuela batholith including the Junin porphyry system, and the Balsapamba‐
Telimbela intrusions), the Calacali‐Pallatanga‐
Pujili fault zone in central Ecuador (associated with the Chaucha batholith), and the diffuse northern limit of the Amotape terrane, probably bracketed between the Piñas‐Portovelo and Jubones faults (associated with the Cangrejos‐
Zaruma intrusive belt; Fig. 2). The locations of single Miocene ore deposits and, especially, ore deposit clusters in Peru are partly 81
82 controlled by intersections of the regional mag‐
matic belt with variably oriented arc‐transverse structures (e.g., Noble & McKee 1999). Southern Ecuador hosts a number of first‐ and second‐
order arc‐transverse structures which represent a structurally favorable environment for porphyry‐
related mineralization (Tosdal & Richards 2001; Richards 2003). Arc‐transverse structures might relate to collision tectonics between the parautochthonous Ecuadorian mainland and the allochthonous forearc block, which undergoes dextral displacement along major fault zones of the Western Cordillera in an oblique subduction setting (Ego et al. 1996; Prodeminca 2000a). Al‐
ternatively, or additionally, transpressional de‐
formation creating arc‐transverse structures in southern Ecuador might partly relate to the post‐
Paleocene 25±12° clockwise block rotation in‐
ferred for the Amotape terrane from paleomag‐
netic studies (Mitouard et al. 1990). The identification of possible arc‐transverse structures in central‐northern Ecuador is ham‐
pered by Quaternary volcano‐sedimentary cover sequences of the Interandean Depression. A pos‐
sible tool for identifying concealed arc‐transverse structures might be the structural correlation of Tertiary intrusions of the Western Cordillera (e.g., the Apuela batholith) with Tertiary intrusions ex‐
posed in the Eastern Cordillera (Aspden et al. 1992), analogous to the structural grain of the Late Oligocene‐Early Miocene Cangrejos‐Zaruma intrusive belt in southern Ecuador, which con‐
nects eastwards with the Paleocene‐Eocene San Lucas pluton (Fig. 2). However, the existing geo‐
chronologic framework for Ecuador’s Eastern Cordillera is largely based on K‐Ar data, which commonly show thermally disturbed Late Creta‐
ceous to Early Tertiary ages (Aspden et al. 1992). Therefore, further U‐Pb zircon geochronologic studies of undeformed Eastern Cordillera intru‐
sions are required to unambiguously confirm their Tertiary age before meaningful structural correlations are possible. Short‐lived compressional Quechua events have been proposed to control the onset or termina‐
tion of Miocene sub‐belt metallogenesis in northern‐central Peru (Noble & McKee 1999; note, however, that some disagreement as to the timing of the different Quechua phases in north‐
ern‐central Peru exists, with different time ranges proposed by Benavides‐Cáceres, 1999, and Noble & McKee, 1999) and might therefore be of similar importance in southern Ecuador. A number of tectonic studies assess the Miocene deformation history of southern Ecuador. The Late Oligocene‐
Early Miocene stress field was probably mostly characterized by horizontal extension, as evi‐
denced by growth sequences of Saraguro Group volcanics forming thickening wedges towards the southern Piñas‐Portovelo fault, indicative of syn‐
volcanic normal fault slip (Spencer et al. 2002). Horizontal extension in southern Ecuador was followed by transpression which is recorded by inversion of the Piñas‐Portovelo fault and folding in the area north of the fault producing a major anticline subparallel to the Cangrejos‐Zaruma intrusive belt (Spencer et al. 2002), as well as by a conjugate set of NE‐trending faults with evidence for dextral movement (Prodeminca 2000a). A compressive pulse at 19 Ma is constrained by the age of Saraguro Group volcanics unconformably overlying deformed sedimentary rocks of the Jacapa Formation in southern Ecuador (Hunger‐
bühler 1997). Furthermore, whole‐scale tilting (30° to the SW) of the Saraguro Group volcanic sequence north of the Piñas‐Portovelo fault is observed (Spencer et al. 2002). Small plutons north of the Piñas‐Portovelo fault which, based on the radiometric age systematics discussed in Chapter 2, can be inferred to be of mainly Early Miocene age, show asymmetric sigmoidal plan‐
view geometries indicative of syntectonic intru‐
sion into a dextral transpressional stress field (Spencer et al. 2002). Geologic evidence thus documents a change from a dominantly tensional to a transpressional stress field in southern Ecua‐
dor in the Early Miocene, which is broadly cor‐
relative in time with the Quechua 1 event of No‐
ble & McKee (1999). Following a period of Mid‐
Miocene extension a second, major compressive Figure 7 (previous page): Geological map of the Gaby and Papa Grande Au‐Cu porphyry systems, and sampling loca‐
tions for geochronology. Samples for U‐Pb zircon datation (Chapter 1) were collected from surface outcrop exposure whereas samples for Re‐Os molybdenite and U‐Pb titanite datation are drill core sample (cf. Tab. 2). Adapted from Sri‐
vastava et al. (2008). 83
84 pulse at c. 9 Ma led to widespread basin inversion in southern Ecuador (Hungerbühler et al. 2002) and is approximately correlative with the Quechua 2 event of Noble & McKee (1999). Ages obtained in this study show that Miocene mineralization in Ecuador clearly predates, partly overlaps with, and postdates the regional “Quechua 1” and “Quechua 2” events (Fig. 10). The “Quechua 3” event seems to terminate Mio‐
cene mineralization in Ecuador. This relationship is rather coincidental, however, as it correlates with the cessation of arc magmatism in southern Ecuador where minimum ages progressively young northwards until the southern end of the active Northern Volcanic Zone in response to slab flattening below southern Ecuador (and northern Peru; Chapter 2). We therefore argue that re‐
gional compressive events do not seem to signifi‐
cantly control Miocene metallogenesis in Ecua‐
dor. Rather, favorable conditions for intrusion emplacement and mineralization at structurally favorable sites (see above) might be associated in time with local stress regime changes, particu‐
larly at the onset of local post‐compressional dila‐
tion (Prodeminca 2000a). The connectivity between the Mio‐
cene metallogenic belts of northern‐
central Peru and Ecuador Sillitoe (1988) popularized the concept of seg‐
mented metallogenic belts in the Andes and placed a fundamental segment boundary at 5°S, corresponding to the Huancabamba Deflection (Fig. 1). Due to a lack of geochronologic data for Ecuadorian ore deposits at that time Sillitoe (1988) noted that the continuity of the northern‐
central Peruvian Miocene metallogenic belt to‐
wards Ecuador is uncertain. Geochronologic data presented in this study as well as geometric con‐
tinuity in map view (Fig. 1) strongly suggest that the Miocene metallogenic belt of northern‐
central Peru is continuous at least until southern Ecuador. Due to the relatively sparse occurrence of Miocene ore deposits in central‐northern Ec‐
uador and southern Colombia, the northward continuation of the belt is less well defined. However, given the continuous distribution of Miocene magmatism along the Ecuadorian mar‐
gin (Chapter 2), and the punctual occurrence of Miocene ore deposits in central‐northern Ecua‐
dor (Junin, Balsapamba, Telimbela), the metal‐
logenic belt might continue further northwards and connect with the metallogenic belt segment of Colombia (Fig. 1; Sillitoe 1988). The typically vertically stacked environments of epithermal vs. porphyry‐style mineralization (e.g., Fontboté & Bendezú 2009) imply that the erosion level at the deposit scale constitutes a major con‐
trol factor for the exposed ore deposit type. This is well exemplified in the western foothills of the Western Cordillera in Ecuador which are deeply eroded and mostly present the cores of Miocene porphyry systems (Prodeminca 2000a). In con‐
trast, large parts of the Azuay district in southern Ecuador are less deeply eroded and preserve abundant Miocene epithermal mineralization (Prodeminca 2000b). Quaternary volcanics of the northern‐central Ecuadorian Interandean Depres‐
sion can be expected to conceal Tertiary arc magmatic units (Chapter 2) and, potentially, Mio‐
cene mineralization at relatively shallow depth. The close geometric similarity between the Peru‐
vian and Ecuadorian belt segments calls for a closer inspection of their respective geologic characteristics which are summarized in Table 5. Both belt segments essentially host ore deposits Figure 8 (previous page): Typical mineralization/alteration assemblages for molybdenite samples dated in this study (red rectangles mark used vein type). A – hornblende granodiorite porphyry with pervasive potassic, overprinted by
phyllic alteration, cut by multiple mo‐qtz veinlets (Junin). B – hbl‐bt granodiorite with weak pervasive chl‐act alteration, vein‐like aggregates of bt‐ep‐py‐cp, and thin qtz‐py‐mo‐cp and mo‐qtz veinlets (Telimbela). C – Vein breccia of grano‐
diorite with weak chl‐ep‐ab‐rt alteration; breccia matrix = ms‐bt‐qtz with disseminated cp‐py‐mo (Balsapamba). D –
Hydrothermal breccia with plag porphyry clasts (with strong ser alteration) and matrix of qtz‐chl‐rt‐goe‐sulfides (po‐cp‐
mo); ser (phyllic) alteration seems to postdate brecciation such that the breccia matrix mineralogy might not be original (Gaby). E – Tonalite with strong chl‐ser/ms alteration and hydrothermal rt (possibly residual from earlier potassic al‐
teration?); cut by multiple veinlets of qtz‐py (with ser halo), mt‐cp‐qtz‐mo, and qtz‐mo (Chaucha). F – Banded cp‐mo‐
jsp‐sl veins with pervasive clay alteration (Tres Chorreras). G – Jigsaw breccia with angular, silicified volcanic clasts in jsp‐tm‐mo‐mt matrix (Tres Chorreras). Scale bar is 2 cm. Same mineral abbreviations as in Fig. 3. 85
of the same types and age range with the possi‐
ble exception of skarn and Cordilleran vein type deposits; the latter, although very common in Peru (Noble & McKee 1999; Fontboté & Bendezú 2009), are not clearly described as such in Ecua‐
dor. This difference might only be an apparent one, caused by different ore deposit classification standards in Ecuador, and/or could be related to the scarcity of reactive limestone host units in Ecuador compared to their abundant occurrence in Peru (e.g., the Pucará Group; Noble & McKee 1999) where they host economically important polymetallic replacement bodies (Fontboté & Bendezú 2009). Contrasting deep crustal base‐
ment compositions in Ecuador and Peru (oceanic vs. continental; Tab. 5) do not seem to produce major differences in metallogenesis between these arc segments, although crustal contribu‐
tions may influence commodity proportions of porphyry Cu deposits elsewhere (e.g., Mo; See‐
dorff et al. 2005). The temporal distribution pattern of the north‐
ern‐central Peruvian Miocene metallogenic belt shows two maxima for radiometric mineralization ages at 15‐13 Ma and 10‐7 Ma (Noble & McKee 1999; note that these maxima are not weighted by tonnage). A few Ecuadorian porphyry Cu de‐
posits such as Chaucha (c. 9.5‐9.9 Ma) and Junin (several events between 9 and 6 Ma) overlap in age with the younger age peak identified in Peru, and polymetallic mineralization at Tres Chorreras (12.9‐12.8 Ma) and advanced argillic alteration at El Mozo (15.4 Ma; Prodeminca 2000b) tend to overlap with the older age peak in Peru. In gen‐
eral, however, data for Ecuador are too scarce to allow a representative statistical treatment at this point. Qualitatively, deposit formation in the Ecuadorian belt segment additionally tends to peak in the Early Miocene, comprising the por‐
phyry Cu deposits of Telimbela, Balsapamba, Gaby‐Papa Grande, and Cangrejos; there is no equivalent peak in the northern‐central Peruvian belt segment at that time. One might speculate that this age difference partly relates to a sam‐
pling bias due to differential exposure levels in various ore deposit districts. As noted above, Early Miocene ore deposits in the Western Cordillera and in the Cangrejos‐
Zaruma intrusive belt of Ecuador represent rela‐
tively deeply eroded porphyry systems domi‐
nated by sodic‐calcic and potassic alteration zo‐
nes (e.g., Prodeminca 2000a; Fig. 3). In contrast, the younger Ecuadorian porphyry deposits such as Chaucha and Junin consistently display wide‐
spread phyllic alteration zones indicating that they have been less deeply eroded. Spikings et al. (2005) show that increased Mid‐ to Late
Figure 9: Concordia plot and weighted mean 206Pb/238U average age for hydrothermal titanite related to sodic‐calcic alteration at the Gaby porphyry system. 86 Miocene cooling rates inferred from thermo‐
chronologic modeling relate to the rapid exhuma‐
tion of fault blocks in the Western Cordillera of Ecuador. Similarly, N‐S contraction (see discus‐
sion above) might have driven Mid‐ to Late Mio‐
cene exhumation of the Cangrejos‐Zaruma intru‐
sive belt which intrudes the hinge and southern flank of a regional antiform north of the inverted Piñas‐Portovelo fault (Prodeminca 2000a; Spencer et al. 2002). It might thus be speculated that younger ore deposits in deeply eroded blocks were completely removed by erosion; a higher density of geochronologic studies on Ec‐
uadorian ore deposits in less deeply eroded fault blocks might preferentially reveal younger min‐
eralization ages. This includes, for example, a large number of potentially younger epithermal deposits in the Azuay district of Ecuador (e.g., Prodeminca 2000b). Geodynamic impacts on the spatio‐
temporal distribution of Miocene ore deposits in Ecuador Two geodynamic factors have been proposed to influence the formation of porphyry‐related ore deposit in the Andes and elsewhere: slab shal‐
lowing‐flattening (e.g., James & Sacks 1999; Kay et al., 1999; see also review by Richards, 2003, and references therein) and the subduction of bathymetric highs, i.e. seamount chains (“ridges”; e.g., Rosenbaum et al. 2005; Cooke et al. 2005). As they represent overthickened oceanic crust, the subduction of seamount chains may contrib‐
ute to slab shallowing by virtue of their relative buoyancy, thus mutually linking ridge collision and flat subduction (e.g., van Hunen et al. 2004). Slab dehydration during periods of flat subduc‐
tion may lead to intense hydration of the overly‐
ing crust; subsequent slab re‐steepening and trenchward migration of hot asthenosphere leaves the thus rheologically weakened crust highly susceptible to deformation and melting which is potentially favorable for metallogenesis (James & Sacks 1999). Shallowing of the slab principally results in the landward migration of the locus of arc magmatism on the upper plate, and may result in porphyry‐related ore deposit formation at the end of a given tectono‐
magmatic cycle under favorable exposure levels (Kay et al. 1999; Richards 2003). The distribution pattern of Tertiary arc magma‐
tism in northern‐central Ecuador suggests a broadly stable slab dip until the Late Miocene when minor slab shallowing occurred, possibly related to subduction of the Carnegie Ridge sea‐
mount chain (Chapter 2). Slab shallowing might thus contribute to generating a favorable expo‐
sure level for young ore deposits in northern Ec‐
uador (such as Junin) by causing an eastward mi‐
gration of Late Miocene to Quaternary arc vol‐
canism. Similarly, progressive broadening of the flat slab region below southern Ecuador (and northern Peru) since the Mid‐Miocene generated a favorable exposure level for porphyry‐related ore deposits in this arc segment where post‐Late Miocene volcanic cover sequences are lacking (Chapter 2). Rosenbaum et al. (2005) present geodynamic reconstructions constraining the collisional timing of anomalous oceanic bathymetric features, the Inca plateau and the Nazca ridge, with respect to the Peruvian margin. These authors observe an apparent spatio‐temporal coincidence between bathymetric high collision and Miocene ore de‐
posit formation (note, however, that these au‐
thors do not quantify any uncertainties for their reconstruction models), and propose that a direct link exists between the two. In this context it is interesting to investigate the metallogenic re‐
sponse to collision of the Carnegie Ridge with the Colombian‐Ecuadorian margin. The timing of initial collision of the Carnegie Ridge with the Colombian‐Ecuadorian margin cannot be accurately determined because, unlike for other ridges such as the Nazca Ridge, a sym‐
metric mirror hotspot track of the Carnegie Ridge does not exist. While the Cocos Ridge was formed at the same time at the Galapagos hotspot as the Carnegie Ridge, it does not represent a direct mirror image due to repeated jumps of the Nazca‐Cocos spreading center across the hotspot position (Barckhausen et al. 2008). Furthermore, the leading edge of both ridges has been sub‐
ducted (e.g., Gutscher et al. 1999). Consequently, the shape of the subducted part of the Carnegie 87
Figure 10: Spatio‐temporal distribution of Miocene ore deposits and associated Oligocene‐Miocene plutons of Ecuador. Mineralization essentially comprises the whole Miocene (and the latest Oligocene: Cangrejos porphyry, 23.5±0.1 Ma)
and seems to peak in the Early Miocene, with a second, broader peak in the Mid‐ to Late Miocene. A general spatio‐
temporal correlation between ore deposit formation and Carnegie Ridge subduction or compressive pulses is not ob‐
served. In a given batholith system (Apuela, Balsapamba‐Telimbela, Chaucha) preserved porphyry‐related mineraliza‐
tion tends to occur after a significant lag time with respect to initiation of batholith construction (c. 5‐15 m.y.); miner‐
alization is concomitant with the last pulse of magmatism and final batholith assembly for Apuela and Chaucha, but
followed by minor ongoing magmatism at Balsapamba‐Telimbela (though note that the latter observation is solely
based on K‐Ar ages which are susceptible to disturbance by younger hydrothermal alteration events). Mineraliza‐
tion/alteration ages compiled from this study or references discussed in the text; magmatic ages from compilation in Chapter 2. Note that only intrusions spatially associated with mineralization are shown. Compressive pulses in Ecuador
(black boxes; Hungerbühler et al. 2002 and references therein), and northern‐central Peru (gray boxes: Noble & McKee 1999; white boxes: Benavides‐Cáceres 1999) are shown for comparison (I = Inca; Q = Quechua). Note that minimum
size of boxes, corresponding to 1 m.y., was arbitrarily chosen and does not reflect the actual duration of compressive event. See text for further discussion. 88 Table 5: Comparison of geological features of the Miocene metallogenic belts of Peru and Ecuador
northern-central Peru
Ecuador
Main commodities
Mineralization age range
Mineralization peaks
Ore deposit types
Au, Cu, base metals
23 - 6 Ma
15 - 13 Ma; 8 - 7 Ma
porphyry Au-Cu, porphyry Cu-Mo,
high-sulfidation epithermal, Cordilleran vein, skarn
Au, Cu
24 - 6 Ma
not enough data available to evaluate
porphyry Au-Cu, porphyry Cu-Mo, high-,
intermediate-, low-sulfidation epithermal
Main host lithologies of ore
deposits
Mesozoic shelf carbonates and
sediments
Late Cretaceous-Paleogene island arc volcanics (northern Ecuador) and oceanic plateau units (central Ecuador)
Neogene volcanics-intrusions
Paleozoic-Mesozoic metasediments (southern Ecuador)
Neogene volcanics-intrusions
(southern Ecuador)
Deep crustal basement units
Mature continental crust
Oceanic plateau crust (Western Cordillera)
Oceanic ± continental crust (southern Ecuadorian Sierra)
Continental crust (Cangrejos-Zaruma intrusive belt)
Data for Ecuador from this study and Prodeminca (2000a, 2000b). Data for Peru from Noble & McKee (1999) and Noble et al.
(2004).
Note that the age of the oldest deposit dated in this study (Cangrejos: 23.5 Ma) corresponds to the latest Oligocene
Ridge is not known, although it seems to be visi‐
ble in seismicity distribution patterns by causing a seismic gap (Gutscher et al. 1999). Predicting the initial collision of the Carnegie Ridge with the Ecuadorian margin thus requires a kinematic reconstruction of the ridge track through time assuming a fixed origin at the Gala‐
pagos hotspot, followed by progressive stepwise rotation using rotation poles for a given plate tec‐
tonic reference system. Previous reconstructions in this manner obtained initial collision estimates around 8 Ma (using rotation poles of Pilger 1984; e.g., Daly 1989; Gutscher et al. 1999), although Lonsdale & Klitgord (1978) estimated a more re‐
cent collision at c. 1 Ma. Using minimum/ maxi‐
mum plate convergence rates of Pardo‐Casas & Molnar (1987) instead of rotation pole data, Spik‐
ings et al. (2001) estimated an initial collision of the Carnegie Ridge with the Ecuadorian margin at 9 or 15 Ma, respectively, for a starting reference time of 22 Ma. We have performed a kinematic Carnegie Ridge reconstruction using the same starting reference time (22 Ma) as Spikings et al. (2001), and the most recent sets of available rotation poles of the Nazca‐South America reference system (Somoza 1998) and of the hotspot‐South Amer‐
ica/Farallon/Nazca reference systems (Müller et al. 1993; Rosenbaum et al. 2005). This approach enables us to reconstruct the ridge assuming ei‐
ther a fixed hot spot, or a fixed South America as reference. Results are presented as a series of time slices in Figure 11 and yield initial collision estimates of 8±3.5 Ma (hotspot‐Farallon/Nazca and Farallon/Nazca‐South America reference sys‐
tems) and 5±4.5 Ma (hotspot‐Farallon/Nazca and hotspot‐South America reference systems). Er‐
rors include a 25% relative uncertainty estimate for the angular velocity of a given rotation pole (cf. Chapter 2) and a ±50 km uncertainty for the hotspot location. Both ages overlap within error and are in good agreement with previous propo‐
sitions of initial Carnegie Ridge collision at c. 8 Ma. It should be noted, however, that (1) these calculations (like all plate tectonic reconstruc‐
tions) are based on the assumption of a plate be‐
having as a rigid entity which might not be strictly true at plate margins (e.g., Cox & Hart 1986), and (2) the reconstructions model a linear hotspot track whereas spreading center jumps (Barck‐
hausen et al. 2008) might have caused ridge seg‐
mentation.
89
Figure 11: Total reconstructions of the position of a reference point originating at coordinates of the central Galapagos
hotspot at different times (boxes in the upper left) throughout the Miocene using the methods described in Cox & Hart (1986). The reference point serves as a proxy for the leading edge of the offshore Carnegie Ridge seamount chain. Two different reference frames were used: a fixed South American plate and the relative motions of hotspots and the
Nazca/Farallon plate (Müller et al. 1993; Somoza 1998), or a fixed hotspot and the relative motions of South America and the Nazca/Farallon plate (Müller et al. 1993; Rosenbaum et al. 2005). The different reference frames constrain ini‐
tial Carnegie Ridge collision at 8±3.5 Ma or 5±4.5 Ma, respectively. Note that in the fixed South America‐based recon‐
struction, reference point positions are theoretical and only become geologically relevant at the trench where the
Nazca/Farallon and South American plates are in direct contact. In contrast, for fixed hotspot‐based reconstructions offshore locations have geological relevance. Carnegie Ridge seamounts dated by Christie et al. (1992) were rotated
back to their origin to test rotation pole accuracy; they mostly overlap with the hotspot starting position suggesting that the former are accurately estimated. 90 Figure 11 (continued) The time‐space distribution of Ecuadorian por‐
phyry‐related ore deposits shows that there is no spatio‐temporal coincidence between ore deposit formation and Carnegie ridge collision, except for the Junin porphyry Cu‐Mo deposit (Fig. 10). Ore deposits in southern Ecuador also predate the arrival of the inferred Inca plateau at the Peru‐
Ecuador trench, which Rosenbaum et al. (2005) estimate at c. 14‐12 Ma. Rosenbaum et al. (2005) speculate that ridge collision causes increased deformation in the upper plate which might aid ore deposit formation. As discussed by MacMillan et al. (2004) and Michaud et al. (2009), however, a direct causative relation between ridge‐trench collision and geological features (such as defor‐
mation) of the overriding plate usually cannot be demonstrated unambiguously. An exception is the forearc region where, depending on its
91
rheological strength, ridge collision either induces uplift, or indentation and conjugated strike‐slip faulting (Hampel et al. 2004). Cooke et al. (2005) note that ridge collision (and flat subduction) is no requirement for the formation of average‐size porphyry‐related ore deposits, but might consti‐
tute a positive trigger mechanism for the forma‐
tion of giant ore deposits. More detailed studies exploring the potential connection between Car‐
negie Ridge‐margin collision, tectonomagmatic processes, and porphyry‐related mineralization at Junin might contribute to unravel these mecha‐
nisms in more detail. Conclusions We have presented the first regionally extensive dataset of radiometric ages on Miocene ore de‐
posits of Ecuador based on the Re‐Os (molyb‐
denite) and U‐Pb (titanite) isotopic systems. These new data allow us to infer that the Mio‐
cene metallogenic belt of northern‐central Peru extends northwards into southern Ecuador, and potentially further north until Colombia. The ages of intrusions and their related hydrothermal sys‐
tems in Ecuador coincide with the age range of porphyry‐related ore deposits in northern‐central Peru and range from 6 to 23.5 Ma. Miocene porphyry‐related mineralization is rela‐
tively widespread in southern Ecuador, and facili‐
tated by metallogenetically favorable factors in‐
cluding (1) a structural setting providing abun‐
dant arc‐transverse structures to channelize magmas and fluids; (2) widespread Miocene magmatism and synmagmatic deformation; (3) erosion levels suitable to preserve Miocene min‐
eralization; and (4) lack of extensive Quaternary cover sequences in a flat slab segment. Miocene porphyry systems in the Western Cordillera and its western foothills may be deeply eroded such that the bulk of the ore body is eroded, but are well‐preserved in some fault blocks. Porphyry‐related ore deposits in Ecuador are of‐
ten (at Junin, Telimbela, Balsapamba, and Chaucha) associated with variably long‐lived magmatic cycles of batholith construction where batholithic precursor magmatism has lifetimes of 5‐15 m.y. Batholithic complexes are associated with elevated transcrustal melt flux and accumu‐
lation, and may thus represent potentially favor‐
92 able sites for porphyry‐related mineralization. However, extensive shallow crustal magmatism associated with peaks in batholith construction (potentially including catastrophic voluminous ignimbrite eruptions) might prove disadvanta‐
geous from a metallogenic perspective. In con‐
trast, during the waning stages of batholith as‐
sembly when thermal relaxation occurs and the focus of magmatism migrates from upper to‐
wards deeper crustal levels favorable petroge‐
netic preconditioning of potential porphyry pa‐
rental melts for subsequent porphyry‐related mineralization might take place. A general direct relationship in space and time between seamount chain subduction or pulses of regional compression and ore deposit formation is not observed in Ecuador except, perhaps, for the Junin porphyry Cu‐Mo deposit in NW Ecua‐
dor. The subduction of bathymetric anomalies such as the Carnegie Ridge or the Inca Plateau may, however, have a metallogenetically favor‐
able effect by their influence on slab dip (with a potential lag time of several m.y.) and thus sur‐
face migration patterns of arc magmatism culmi‐
nating in an arc magmatic gap. The latter results in favorable exposure levels of Miocene por‐
phyry‐related ore deposits in southern Ecuador and northern Peru. References Aspden, J.A., Harrison, S. H., Rundle, C. C. (1992): New geochronological control for the tectono‐magmatic evolution of the metamorphic basement, Cordillera Real, and El Oro Prov‐ince of Ecuador. Journal of South American Earth Sciences 6; 77‐96. Bachmann, O., Miller, C. F., de Silva, S. L. (2007): The volcanic‐plutonic connection as a stage for under‐
standing crustal magmatism. Journal of Volcanology and Geothermal Re‐search 167; 1‐23. Barckhausen, U., Ranero, C R., Cande, S. C., Engels, M., Weinrebe, W. (2008): Birth of an intraoceanic spread‐
ing center. Geology 36; 767‐770. Barra, F., Ruiz, J., Mathur, R., Titley, S. (2003): A Re‐Os study on sulfide minerals from the Bagdad porphyry Cu‐Mo deposit, northern Arizona, United States. Min‐
eralium Deposita 38; 585–596. Barra, F., Ruiz, J., Valencia, V. A., Ochoa‐Landín, L., Chesley, J. T., Zurcher, L. (2005): Laramide porphyry Cu‐Mo mineralization in northern Mexico: age con‐
straints from Re‐Os geochronology in molybdenite. Economic Geology 100; 1605‐1616. Benavides‐Cáceres, V. (1999): Orogenic evolution of the Peruvian Andes: The Andean cycle. In: Skinner, B. J. (ed.), Geology and ore deposits of the central Andes. SEG Special Publication 7; 61–107. Bineli Betsi, T. (2007): The low‐sulfidation Au‐Ag de‐
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tonics 25; TC3017. 96 CHAPTER IV CRUSTAL BASEMENT ARCHITECTURE IN ECUADOR EXPLORED BY Sr, Nd, AND Pb ISOTOPIC COMPOSITIONS OF TERTIARY‐QUATERNARY ARC MAGMAS Abstract The crustal basement of Ecuador comprises a collage of mostly Paleozoic‐Mesozoic tectono‐stratigraphic units of both continental and oceanic affinity in the Eastern Cordillera, and oceanic plateau units in the Western Cordillera and forearc region which were accreted in the Late Cretaceous. The diffuse paleo‐
continental suture zone is bracketed by the regional Andean‐ (NNE) trending Peltetec and Calacali‐
Pallatanga‐Pujili fault zones, and is situated between the Eastern and Western Cordillera ranges where basement units are covered by Tertiary‐Quaternary arc volcanics. An extensive body of isotope geochemi‐
cal information exists for Quaternary arc volcanics of the Northern Volcanic Zone in northern Ecuador, whe‐
reas Tertiary cover sequences in southern‐central Ecuador are poorly characterized isotopically. In this study we are presenting a set of 58 new whole‐rock Sr, Nd, and Pb isotopic compositions of Late Oligocene and younger intrusions and arc volcanics of the Western Cordillera, its western foothills, and the central‐
southern Ecuadorian Sierra region. Combining this new dataset with existing data on Quaternary arc volca‐
noes allows us to trace basement units of the Late Cretaceous suture zone at depth. Quaternary arc volcanics define distinct isotopic groups for volcanoes situated east and west of the Peltetec Fault, respectively. Late Oligocene to Late Miocene arc volcanics and intrusions of the southern Ecuador Sierra region overlap isotopically with recent arc volcanics east of the Peltetec Fault in northern Ecuador suggesting along‐strike continuity of similar basement units at depth. Late Oligocene to Late Miocene grani‐
toids of the Western Cordillera and its western foothills show the most primitive Sr and Nd isotopic compo‐
sitions identified in Tertiary‐Quaternary Ecuadorian arc magmas so far; primitive Cretaceous oceanic pla‐
teau units constitute their assimilants at depth, causing these arc magmas to become isotopically more primitive while assimilating crustal material. Crustal isotopic contamination of Tertiary‐Quaternary arc magmas mainly takes place at deep to mid‐crustal levels except for granitoids of the Cangrejos‐Zaruma in‐
trusive belt in southern Ecuador, where additional prominent shallow crustal assimilation produces highly radiogenic Sr and Pb, and low radiogenic Nd isotopic compositions of evolving arc magmas. Isotopic Sr, Nd, and Pb compositions of arc magmas in northern‐central Ecuador follow a systematic across‐
arc pattern where they evolve towards progressively more radiogenic 87Sr/86Sr and 207Pb/204Pb, and less ra‐
diogenic 143Nd/144Nd compositions at deep to mid‐crustal levels with increasing distance from the trench. This is consistent with regional underthrusting of accreted oceanic plateau material along a broad suture zone below the paleo‐continental margin as previously inferred from seismic studies. 97
Introduction Geochemical studies in the southern and central Andes have revealed pronounced variations in Sr, Nd, and Pb isotopic compositions of arc magmas. These variations are mainly attributed to crustal contamination effects where lateral and vertical basement heterogeneity (Davidson & de Silva 1992; Wörner et al. 1992), crustal thickness (Hil‐
dreth & Moorbath 1988), and the extent of direct magma‐crust interaction (Dungan & Davidson 2004) are inferred to be the major control factors for the isotopic variability of arc magmas. Crustal isotopic imprints on arc magmas may be super‐
posed on continent‐scale mantle wedge isotopic variability (Chiaradia & Fontboté 2002) which might, for example, involve source contamination by subduction erosion (e.g., Stern & Skewes 2005). In a given arc segment, the isotopic char‐
acteristics of Tertiary‐Quaternary arc magmas may be used to outline different crustal base‐
ment domains (e.g., Wörner et al. 1992; Mamani et al. 2008, 2010). In the northern Andes, significant crustal base‐
ment variations occur both in along‐ and across‐
arc dimension. The Tertiary‐Quaternary arc sys‐
tem of Ecuador is constructed on a basement collage of multiple tectono‐stratigraphic units separated by major NNE‐trending fault zones (Li‐
therland et al. 1994). Juxtaposed against the cra‐
tonic basement of the Amazon foreland basin, several Paleo‐ to Mesozoic continental and island arc units form the major basement units of the Eastern Cordillera (Fig. 1; Litherland et al. 1994), alternatively interpreted as autochthonous conti‐
nental margin (Pratt et al. 2005). Exotic Creta‐
ceous oceanic plateau fragments floor the West‐
ern Cordillera and forearc region (e.g., Mamberti et al. 2003; Vallejo et al. 2009). Geophysical stud‐
ies yield contradictory results with respect to the geometry of the boundary between the accreted oceanic plateau material and the parautochtho‐
nous paleocontinental domain below the Inter‐
andean Depression (IAD) between Ecuador’s Eastern and Western Cordilleras. Based on gra‐
vimetric data Feininger & Seguin (1983) suggest that continental Eastern Cordillera basement floors the IAD; in contrast, Guillier et al. (2001) provide seismic evidence for regional‐scale un‐
derthrusting of oceanic plateau material below the IAD. While isotopic data for individual arc volcanic centers and the accreted oceanic domains are readily available (see references in Figs. 1, 2), on‐
ly a single study has attempted to comprehen‐
sively assess the Pb isotopic composition of the different terrane basement units on a regional scale (Chiaradia et al. 2004a). In this contribution, we are presenting 58 new whole‐rock Sr, Nd, and Pb isotopic compositions of Late Oligocene and younger granitoids and volcanic formations (and some of their host lithologies) of the Western Cordillera, its western foothills, and the central‐ southern Ecuadorian Sierra region. We are com‐
bining these new data with published isotopic compositions of basement units and Quaternary arc volcanoes to discuss tectonic implications of isotopic variations in the crustal basement of the IAD and the southern Ecuadorian Sierra, and how Figure 1 (next page): Left: Topographic map of northern Andean margin showing gravity anomaly isolines [mgal] of Feininger & Seguin (1983) in the Cordillera region (thick white lines). Note the overall lower elevation in southern
Ecuador‐northern Peru compared to central‐northern Ecuador where oceanic plateau units are underthrusting the continental margin. Right: Simplified geological map of the Ecuadorian Andes showing crustal basement units, major fault systems (straight lines; dashed where inferred; adapted from Winkler et al. 2005), Late Cretaceous‐Tertiary arc volcanic units and intrusions, and Quaternary arc volcanoes relevant for this study. New Sr, Nd, and Pb isotopic data
presented in this study were obtained on Late Tertiary granitoids and volcanics (and the Late Cretaceous Curiplaya intrusions). Arc volcanoes and intrusions are color‐coded according to their isotopic composition reflecting crustal
basement domains as discussed in the text (yellow: εNdinitial >5, 87Sr/86Sr <0.7038, 206Pb/204Pb <18.9, 207Pb/204Pb <15.62; blue: εNdinitial = 4‐6, 87Sr/86Sr = 0.7038‐0.7044, 206Pb/204Pb <19.03, 207Pb/204Pb <15.60; green: εNdinitial = 3‐6, 87
Sr/86Sr = 0.7040‐0.7043, 206Pb/204Pb < 19.14, 207Pb/204Pb < 15.64; brown: εNdinitial > ‐2, 87Sr/86Sr < 0.7049, 206Pb/204Pb <19.08, 207Pb/204Pb <15.70; dark‐gray: εNdinitial <1, 87Sr/86Sr >0.7047, 206Pb/204Pb >18.9, 207Pb/204Pb >15.62). Volcanoes where only a single isotopic analysis is available (Quilotoa, Sangay, Chalupas) are marked with a question mark and
not color‐coded unless the isotopic signature is unambiguous. Note the systematic across‐arc distribution pattern of arc magma isotopic compositions in northern Ecuador as discussed in the text. General geological features of map modified from Chapter 2. 98 99
they relate to the proposed suture zone geome‐
tries of oceanic plateau units and the paleoconti‐
nental Ecuadorian margin. No attempt is being made here to extend the discussion towards iso‐
topically discerning the complex Eastern Cordil‐
lera basement (Litherland et al. 1994; Pratt et al. 2005). Geological framework Paleozoic and subordinate Precambrian base‐
ment units of Ecuador’s Eastern Cordillera host a number of major batholiths and volcanics result‐
ing from intense Triassic‐Jurassic arc magmatism (Litherland et al. 1994). While Litherland et al. (1994) tend to interpret major batholith bounda‐
ries as suture zones delineating a number of both oceanic‐ and continental‐affinity allochthonous terranes (Fig. 1), Pratt et al. (2005) regard most of these contacts as intrusive and infer an autoch‐
thonous crustal basement for Ecuador’s Eastern Cordillera. Overthickened oceanic crust, partly associated with the Caribbean‐Colombian oce‐
anic plateau (CCOP) and mainly accreted in the Late Cretaceous, is juxtaposed against the East‐
ern Cordillera basement along a suture zone comprising parts of the IAD (e.g., Mamberti et al. 2003; Jaillard et al. 2005; Spikings et al. 2005; Val‐
lejo et al. 2009). The allochthonous oceanic pla‐
teau units host several pre‐ and post‐accretionary island arc systems of Late Cretaceous to Early Tertiary age (Rio Cala and Macuchi units; e.g., Chiaradia 2009; Vallejo et al. 2009). Mid‐ to Late Tertiary arc magmatism in northern‐central Ec‐
uador, resulting from the subduction of the Faral‐
lon/Nazca plate below the accreted oceanic pla‐
teau material, focused on the Western Cordillera region and only sporadically affected the Eastern Cordillera until the Late Miocene, when major landwards arc broadening towards the Eastern Cordillera is recorded (Chapter 2). Active arc vol‐
canism of the Northern Volcanic Zone (NVZ) in Ecuador covers the whole range from the Eastern to the Western Cordillera, and from the Colom‐
bian border until Sangay volcano at c. 2° S (Fig. 1). Tertiary‐Quaternary volcano‐sedimentary cover sequences conceal the basement of the IAD and the central Ecuadorian Sierra region between the 100 Eastern and Western Cordillera ranges. Major arc‐parallel fault systems thought to bracket the diffuse suture zone between oceanic plateau units and the paleo‐continental margin below the IAD (Calacali‐Pallatanga‐Pujili fault zone, CPPF, and Peltetec fault, PF; Winkler et al. 2005) tend to dip subvertically at the surface, but define c. 35°E dipping fault planes at mid‐ to deep crustal levels implying that oceanic material forms the deep crustal root of the IAD (Guillier et al. 2001; Jaillard et al. 2005); a tectonized mélange of con‐
tinental crust and oceanic plateau units is in‐
ferred at shallow crustal depth (e.g., Spikings et al. 2005). In contrast, IAD basement units similar to the parautochthonous Eastern Cordillera have been inferred in previous studies (Chaucha ter‐
rane; Feininger & Seguin 1983; Litherland et al. 1994). Seismic studies constrain the crustal thickness in Ecuador to 40‐50 km below the Western Cordil‐
lera in the present‐day frontal arc, and to 50‐75 km below the IAD and the Eastern Cordillera in the present‐day main arc region (Guillier et al. 2001). The significantly lower mean elevation of the Ecuadorian Andes compared to the Central Andean Altiplano region of only slightly higher crustal thickness might be isostatically supported by a column of continental crust underthrusted by high‐density oceanic plateau material, follow‐
ing the inferred suture zone geometry at depth (Guillier et al. 2001). Periods of crustal thickening in Ecuador are not as well constrained as in the southern‐central Andes; tectonic crustal thicken‐
ing by westwards basal forearc wedging seems to have occurred throughout the Tertiary with a ma‐
jor period of thickening affecting the Andean main arc region since the Late Miocene (Jaillard et al. 2005). Methodology Samples for isotopic analysis were collected from outcrop exposures or drill cores in Ecuador ac‐
cording to the locations marked in Figure 1. Gra‐
nitoids and arc volcanic formations investigated in this study are of Late Oligocene or younger age (Fig. 1; Chapter 2). Intrusive rocks comprise hornblende‐ and biotite‐bearing tonalites, grano‐
diorites, quartz‐diorites, and granodiorite por‐
phyries emplaced along major fault zones of the Western Cordillera, its western foothills, and in the Sierra region of central‐southern Ecuador. Sampled volcanic rocks of the IAD and the south‐
ern Ecuadorian Sierra comprise (1) the wide‐
spread Late Oligocene to Early Miocene Saraguro Group, composed of andesitic‐dacitic lava flows and tuffs, and dacitic‐rhyolitic ignimbrites; and (2) the Late Miocene Quimsacocha volcanic cen‐
ter which forms a caldera with associated ande‐
sitic‐dacitic lava flows and dacitic‐rhyolitic domes. In addition, we compiled isotopic data of NVZ volcanic centers which, from north to south, comprise Imbabura, Cayambe, Pululagua, Pichin‐
cha, Chacana, Ilalo, Atacazo, Antisana, Cotopaxi, Chalupas, Quilotoa, and Sangay volcano. Sumaco volcano, occupying a back‐arc position 380 km from the trench, was not included in the compila‐
tion because its magmas are significantly en‐
riched in Sr and Nd, minimizing the isotopic lev‐
erage of potential crustal assimilants (Chiaradia et al. 2009). All sample preparation steps and isotopic analy‐
ses were performed at the Department of Miner‐
alogy, University of Geneva. Samples were cleaned with water, crushed using a steel jaw crusher, and powdered (<70 μm) using an agate disc mill. Preparation for isotopic analysis used the techniques of Chiaradia (2009) and refer‐
ences therein. Powdered samples (100‐150 mg each) were loaded into screw‐sealed 20 ml Teflon vials and leached overnight at room temperature using 3 M HCl to dissolve alteration minerals such as carbonates with potential isotopic disequilib‐
rium compositions. Sample‐leachate mixture cen‐
trifugation and subsequent leachate discarding was followed by two‐fold sample residue rinsing and centrifugation using deionized water. Sam‐
ples were then dissolved in a mixture of 4 ml con‐
centrated HF and 1 ml 15 M HNO3 at 140°C for seven days, followed by sample drying on a hot plate, re‐dissolution in 3 ml 15 M HNO3 at 140°C for three days, and a final drying step on a hot plate. Strontium and Nd separation was carried out using cascade columns with Sr‐spec, TRU‐
spec and Ln‐spec resins following a modified method after Pin et al. (1994). Lead was further purified with an AG‐MP1‐M anion exchange resin in hydrobromic medium. Lead, Sr and Nd isotope ratios were measured on a Thermo TRITON mass spectrometer on Faraday cups in static mode. Lead was loaded on Re fila‐
ments using the silica gel technique and all sam‐
ples (and standards) were measured at a pyrome‐
ter‐controlled temperature of 1220 °C. Lead iso‐
tope ratios were corrected for instrumental frac‐
tionation by a factor of 0.07% per amu based on more than 90 measurements of the SRM981 standard and using the standard values of Todt et al. (1996). External reproducibilities (2σ) of the standard ratios are 0.05% for 206Pb/204Pb, 0.08% for 207Pb/204Pb and 0.10% for 208Pb/204Pb. Stron‐
tium was loaded on single Re filaments with a Ta oxide solution and measured at a pyrometer‐
controlled temperature of 1490 °C. 87Sr/86Sr val‐
ues were internally corrected for fractionation using a 88Sr/86Sr value of 8.375209. Raw values were further corrected for external fractionation by a value of +0.03‰, determined by repeated measurements of the SRM987 standard (87Sr/86Sr=0.710250). External reproducibility (2σ) of the SRM987 standard is 7 ppm. Neodymium was loaded with 1 M HNO3 and measured with the double filament technique. 143Nd/144Nd val‐
ues were internally corrected for fractionation using a 146Nd/144Nd value of 0.7219 and the 144Sm interference on 144Nd was monitored on the mass 147
Sm and corrected using a 144Sm/147Sm value of 0.206700. The external reproducibility (2σ) of the JNdi‐1 standard (Tanaka et al. 2000) is 4 ppm. Part of the Nd isotope ratios were measured on a seven‐collector Finnigan MAT 262 thermal ioniza‐
tion mass spectrometer with extended geometry and stigmatic focusing using double Re filaments where 143Nd/144Nd was measured in a semidy‐
namic mode (quadruple collectors, measurement jumping mode). Neodymium isotopic ratios were age‐corrected and recalculated to initial 143Nd/144Nd values us‐
ing appropriate age estimates (Tab. 1) and sam‐
ple compositions (Chapter 5). Strontium isotopic ratios were not age‐corrected because of signifi‐
cant alteration‐induced modifications of whole rock Rb contents rendering Rb‐based 87Sr/86Sr corrections to hypothetic initial values geologi‐
cally meaningless. Most samples are character‐
ized by Rb/Sr ratios of 0.1 or below, and age 101
Table 1: Whole rock 87Sr/86Sr, 143Nd/144Nd, 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb compositions of investigated magmatic centers in Ecuador
Sample
Magmatic center
87
86
Neogene intrusions
E06209
Apuela (Junin)
E06200
Apuela (Cuellaje)
E06202
Apuela (Cuellaje)
E05127
Junin
E06211
Junin
E07032
Junin
E06127
Balsapamba
E06136
Balsapamba
E06131a Balsapamba
E06132
Balsapamba
E06135
Balsapamba
E06138
Balsapamba
E07045
Telimbela
E06150
Telimbela
E06153
Telimbela
E05070
Gaby
E05078
Gaby
E05083b Gaby
E05088
Gaby
E05090
Gaby/Papa
Grande
E06049
Gaby
E06052
Gaby
E07002
Chaucha: Naranjos
E07005
Chaucha: Tunas
E07001
Chaucha: Naranjos
E07003
Chaucha: Tunas
E07008
Chaucha: Tunas
E06158
Chaucha: Gur-Gur
E06071
Cangrejos
E05-M4
Cangrejos
E06069
Cangrejos
E06070
Cangrejos
E06090
Portovelo
E06112
Portovelo
E06115
Portovelo
E06123
Portovelo
E07023
Zaruma
E07016
El Mozo
E07017
El Mozo
E07020
El Mozo
102 Sr
143
Sr
144
Nd
εNd
εNdin
206
204
Nd
Pb
207
Pb
204
Pb
208
Pb
age
age
Pb
204
Pb
[Ma]
ref.*
0.70365
0.70374
0.70379
0.70376
0.70378
0.70377
0.70366
0.70368
0.70369
0.70369
0.70370
0.70371
0.70366
0.70371
0.70372
0.70422
0.70421
0.70416
0.70440
0.70427
0.51298
0.51295
0.51296
0.51297
0.51295
0.51296
0.51293
0.51297
0.51296
0.51297
0.51294
0.51294
0.51297
0.51299
0.51297
0.51290
0.51289
0.51292
0.51286
0.51288
6.7
6.0
6.3
6.4
6.1
6.3
5.6
6.5
6.3
6.5
6.0
6.0
6.5
6.8
6.5
5.1
5.0
5.5
4.4
4.7
6.8
6.1
6.4
6.5
6.2
6.4
5.8
6.7
6.5
6.7
6.1
6.1
6.7
6.9
6.7
5.2
5.1
5.7
4.5
4.9
18.769
18.754
18.875
18.853
18.856
18.876
18.863
18.793
18.692
18.719
18.716
18.710
18.879
18.823
18.800
18.900
18.760
18.679
18.848
19.043
15.595
15.565
15.594
15.602
15.583
15.601
15.592
15.603
15.583
15.619
15.587
15.589
15.611
15.617
15.604
15.587
15.607
15.598
15.603
15.634
38.442
38.394
38.548
38.533
38.474
38.537
38.510
38.502
38.358
38.457
38.361
38.396
38.572
38.553
38.500
38.570
38.461
38.436
38.564
38.706
15
13
13
9
9
9
22
22
21
21
21
21
26
21
16
20
20
20
20
20
1
2
2
2
2
2
2
2
2
2
2
2
2
3
4
2
2
2
2
2
0.70420
0.70427
0.70441
0.51287
0.51289
0.51270
4.5
5.0
1.2
4.7
5.0
1.3
18.972
18.889
19.017
15.642
15.600
15.648
38.731
38.578
38.781
20
20
10
5
5
2
0.70445
0.70430
0.51269
0.51275
1.0
2.2
1.1
2.3
19.032
18.968
15.636
15.617
38.761
38.667
10
15
2
5
0.70430
0.70434
0.70464
0.70646
0.70546
0.70541
0.70560
0.70549
0.70567
0.70588
0.70474
0.70482
0.70478
0.70488
0.70461
0.51276
0.51274
n/a
0.51253
0.51265
0.51255
0.51253
0.51261
0.51263
0.51266
0.51271
0.51269
0.51267
0.51266
0.51269
2.4
2.0
n/a
-2.2
0.2
-1.7
-2.1
-0.6
-0.3
0.4
1.4
1.1
0.6
0.4
1.0
2.5
2.1
n/a
-2.0
0.4
-1.5
-1.8
-0.4
0.0
0.5
1.5
1.2
0.7
0.5
1.1
19.054
19.012
19.012
19.067
19.033
18.923
19.055
18.955
19.040
18.988
18.888
18.991
18.949
18.972
18.934
15.677
15.630
15.642
15.682
15.660
15.631
15.646
15.661
15.666
15.665
15.643
15.639
15.629
15.669
15.634
38.885
38.724
38.764
38.974
38.894
38.745
38.889
38.823
38.922
38.847
38.716
38.791
38.730
38.858
38.735
15
15
15
26
26
23
23
24
24
21
21
21
16
16
16
2
2
5
2
2
6
6
5
2
5
5
2
2
2
2
Table 1 (continued)
Sample
Magmatic center
87
86
Sr
143
Sr
144
Nd
εNd
εNdin
206
204
Nd
Pb
207
Pb
204
Pb
208
Pb
age
age
Pb
204
Pb
[Ma]
ref.*
Neogene volcanics
E05102 Quimsacocha
E06020 Quimsacocha
E05099 Quimsacocha
E06022 Quimsacocha
E06157 Saraguro at
Chaucha
0.70455
0.70457
0.70440
0.70430
0.70476
0.51269
0.51270
0.51272
0.51272
0.51275
1.0
1.1
1.6
1.6
2.1
1.1
1.2
1.7
1.7
2.2
18.990
18.999
18.960
18.917
n/a
15.642
15.650
15.649
15.582
n/a
38.761
38.783
38.744
38.630
n/a
7
7
7
7
25
2
2
5
5
7
E06166
Saraguro at
Chaucha
0.70707
0.51266
0.4
0.5
19.173
15.690
39.026
25
7
E06010
Saraguro at Canicapa
0.70443
0.51274
1.9
2.2
n/a
n/a
n/a
20
7
E06012
Saraguro at Canicapa
0.70440
n/a
n/a
n/a
18.975
15.665
38.837
20
7
E06082
Saraguro at Portovelo
0.70534
0.51261
-0.5
-0.4
18.990
15.667
38.871
25
7
0.70477
0.51274
2.1
2.3
19.038
15.664
38.858
31
2
0.70412
0.70450
0.70419
0.70420
0.51279
0.51282
0.51287
0.51277
2.9
3.5
4.5
2.6
3.6
4.3
5.1
3.4
18.848
18.962
18.890
18.996
15.665
15.622
15.640
15.643
38.744
38.781
38.680
38.833
92
92
92
92
2
2
2
2
0.70371
0.51298
6.6
6.8
18.550
15.593
38.243
40
5
0.70402
0.51301
7.2
7.4
n/a
n/a
n/a
40
5
E07013 Tres Chorreras
Late Cretaceous intrusions
E07029 Curiplaya
E07027 Curiplaya
E07031 Curiplaya
E07028 Curiplaya
Host rocks
E05123 Macuchi (?) at
Junin
E06148 Macuchi at Balsapamba
E06145
Macuchi at Balsapamba
0.70535
0.51301
7.2
7.4
n/a
n/a
n/a
40
5
E06035
CCOP basalt
0.70367
0.51301
7.3
7.5
19.012
15.631
38.768
90
5
*Age references: 1 - MMAJ/JICA 1992; 2 - this study (Chapter 2); 3 - McCourt et al. 1997; 4 - MMAJ/JICA 1991; 5 - estimated
from field relationships and regional geology; 6 - this study (Chapter 3); 7 - Hungerbühler et al. (2002)
corrections using measured, alteration‐
influenced Rb contents typically affect the 87
Sr/86Sr ratio by ±0.0001 or less; this magnitude is below the uncertainty range relevant to the discussion in this article, and no significant bias of our interpretations based on differences be‐
tween measured and initial 87Sr/86Sr ratios is ex‐
pected. Results Compiled Sr, Nd, and Pb isotopic data allow the distinction of three groups of Quaternary volca‐
noes (group 1: εNdinitial = 4‐6, 87Sr/86Sr = 0.7038‐
0.7044, 206Pb/204Pb < 19.03, 207Pb/204Pb < 15.60; group 2: εNdinitial = 3‐6, 87Sr/86Sr = 0.7040‐0.7043, 206
Pb/204Pb < 19.14, 207Pb/204Pb < 15.64; group 3: εNdinitial > ‐2, 87Sr/86Sr < 0.7049, 206Pb/204Pb < 19.08, 207Pb/204Pb < 15.70), following a systematic across‐arc distribution trend in northern Ecuador (Figs. 1, 2; Appendix Tab. A1). The range of iso‐
topic compositions defined by multiple volcanoes in a given group is homogeneous for about 100 km strike length, corresponding to the whole arc segment for which isotopic data is available. The least radiogenic Sr‐Pb and most radiogenic Nd compositions common to all three groups cor‐
respond to the isotopic composition of the petro‐
logically most primitive samples identified by Chiaradia (2009) and Chiaradia et al. (2009) based on isotope and whole rock chemical correlations, 103
Figure 2: Diagrams of εNdinitial vs. 87Sr/86Sr (plus magnified area), 207Pb/204Pb vs. 206Pb/204Pb (plus magnified area), 87Sr/86Sr vs. 206Pb/204Pb, εNdinitial vs. 206Pb/204Pb, and 208Pb/204Pb vs. 206Pb/204Pb isotopic compositions of Late Tertiary (except for Curiplaya) samples analyzed in this study, and isotopic fields for potential crustal
assimilants and NVZ magmas. Orange X marks the isotopic composition of the most primitive melts identi‐
fied by Chiaradia (2009) and Chiaradia et al. (2009), and serves as a proxy for parental melt isotopic compo‐
sitions of Late Tertiary‐Quaternary arc magmas derived from an isotopically broadly homogeneous mantle
wedge. Individual Quaternary arc volcanoes (see Fig. 1 for references) span a relatively narrow field in Sr
isotopes, and show increased isotopic diversity in Nd and Pb isotopes reflecting an oceanic plateau compo‐
nent as main assimilant for volcanic edifices situated west of the Peltetec fault.
104 suggesting that all magmas originated in an iso‐
topically homogeneous mantle wedge. The iso‐
topic arrays defining the three groups are charac‐
terized by variable increases in radiogenic Sr and Pb, and non‐radiogenic Nd components from mi‐
nor (frontal arc, close to the CPPF) to slightly higher (main arc, west of the PF), and significant (rear main arc, east of the PF); 143Nd/144Nd and 207
Pb/204Pb are the most powerful isotopic dis‐
criminators between the three groups. These iso‐
topic distribution characteristics suggest that mantle‐derived arc magmas acquire their distinct isotopic signatures either by variable degrees of assimilation (or mixing with partial melts) of a crustal component with constant isotopic com‐
position, by assimilation (or mixing with partial melts) of crust of different isotopic composition, or a combination of both. While Quaternary vol‐
canoes situated to the west of the PF completely overlap with the range of isotopic compositions of the Macuchi Unit (Eocene island‐arc; Fig. 2), volcanoes east of the fault are characterized by significantly more continental crust‐like Sr‐Nd‐Pb isotopic compositions. Tertiary samples measured in this study (Tab. 1) define homogenous groups in uranogenic and thorogenic Pb isotope and combined Sr‐Nd‐Pb isotope plots for given intrusive suites (Fig. 2). Late Tertiary granitoids of the Western Cordillera foothills (Apuela‐Junin, Balsapamba‐Telimbela) are significantly less radiogenic in Sr and Pb, and more radiogenic in Nd isotopic compositions than the most primitive Quaternary NVZ samples (εNdinitial > 5, 87Sr/86Sr < 0.7038, 206Pb/204Pb < 18.9, 207
Pb/204Pb < 15.62); they isotopically overlap with the most primitive Macuchi and Rio Cala units (their immediate host rocks), and parts of the CCOP (the regional basement unit at depth; Fig. 2). Given the low thickness of the Macuchi Unit (<2.5 km; Kerr et al. 2005) granitoid parental melts likely assimilated primitive CCOP material while they were differentiating at depth. In addi‐
tion, their very limited isotopic variability is con‐
sistent with assimilation of older (Macuchi?) arc intrusive roots with a similar isotopic composition (Dungan & Davidson 2004). Unlike Rio Cala melts (Chiaradia 2009), granitoid parental melts did not interact with seawater‐altered CCOP lithologies characterized by high radiogenic Sr values. Fur‐
ther south, the CCOP basalt‐hosted Gaby intru‐
sive center shows slightly more radiogenic Sr and Pb, and less radiogenic Nd isotopic compositions than Western Cordillera granitoids to the north, and overlaps completely with isotopic composi‐
tions of NVZ frontal arc volcanoes. Late Tertiary intrusions and volcanic rocks of the southern Ecuadorian Sierra are characterized by variably higher radiogenic Sr and Pb, and lower radiogenic Nd isotopic compositions than West‐
ern Cordillera granitoids. They partly overlap with the isotopic compositions of the Cretaceous Tan‐
gula batholith and a number of minor Paleogene intrusions in southern Ecuador, as well as with rear main arc NVZ magmas east of the PF (group 3 above); they consistently plot between poten‐
tial assimilant end‐member isotopic compositions (CCOP and Eastern Cordillera or Amotape base‐
ment; Fig. 2). Several granitoids and volcanics in southern Ecuador completely overlap isotopically with the isotopic compositions of main arc volca‐
noes west of the PF in northern Ecuador. This includes the Mid‐to Late Miocene Chaucha ba‐
tholith, Mid‐Miocene Saraguro Group volcanics at Cañicapa, and the Late Miocene Quimsacocha volcanic center. The El Mozo intrusions, situated at the limit between the Loja and Alao terranes of Litherland et al. (1994), and Saraguro Group vol‐
canics at Tres Chorreras overlap with this group in Nd and Pb isotopic compositions, but are char‐
acterized by a slightly more radiogenic Sr compo‐
sition, although they are still more primitive than the bulk Earth (Fig. 2). A significant contribution of Loja (highly radiogenic Sr and Pb) or Alao Figure 2 (caption continued from previous page): Quaternary volcanoes define distinct isotopic groups as defined in Fig. 1; Late Tertiary intrusions and volcanics of the southern Ecuadorian Sierra region partly overlap with these com‐
positional groups suggesting assimilation of similar basement units. Isotopic reference fields as follows: CCOP (Mam‐
berti et al. 2003; only Western Cordillera outcrops considered); Rio Cala and Macuchi (Chiaradia 2009); Raspas Com‐
plex (Bosch et al. 2002); Western Cordillera shallow basement (mica schist of Amortegui 2007; single Pichincha xeno‐
lith of Chiaradia et al. 2009); Western Cordillera amphibolites/granulites (amphibolites of Amortegui 2007; Pichincha granulite xenoliths of Chiaradia et al. 2009); Amotape, Loja, Alao (Pb; Chiaradia et al. 2004a); Alao, Loja, Cretaceous‐
Paleogene intrusions (Sr, Pb; Chiaradia et al. 2004b). Error bars for Pb isotopes (± 2σ) shown in lower right corner; Sr and Nd isotope error bars are below symbol size.
105
(highly radiogenic Pb) basement units is thus unlikely for parental melts of the El Mozo intru‐
sions. Rather, along with other southern Sierra intrusions and volcanics, they define an isotopi‐
cally homogeneous group suggesting that their parental melts evolved by assimilation processes of similar crustal basement units. The regional distribution of IAD and southern Si‐
erra basement units is relatively, but not com‐
pletely isotopically homogenous, as evidenced by a Saraguro Group tuff collected in the Chaucha area, which shows the most radiogenic Sr and Pb isotopic compositions of the whole dataset. The latter notion is in agreement with the surface exposure of several metamorphic inliers around Chaucha which might represent Amotape (El Oro) basement fragments (Litherland et al. 1994; Pratt et al. 1997) implying that a definite northern Amotape basement border cannot be accurately drawn and rather corresponds to a northwards‐
extending tectonized zone (Spikings et al. 2005). Further south, plutons of the Cangrejos‐Zaruma intrusive belt compositionally extend towards significantly more radiogenic Sr and Pb, and less radiogenic Nd values (εNdinitial <1, 87Sr/86Sr >0.7047, 206Pb/204Pb >18.9, 207Pb/204Pb >15.62) suggesting they assimilated continental crust‐
dominated basement lithologies in the range of the El Oro massif south of the Jubones fault. The Late Cretaceous Curiplaya intrusive center, hosted by the Celica‐Lancones basin in south‐
ernmost Ecuador, compositionally overlaps with the CCOP‐hosted Gaby intrusions in 87Sr/86Sr and 143
Nd/144Nd, but plots at higher radiogenic Pb iso‐
topic ratios. Discussion Isotopic compositional changes in the Ecuadorian mantle wedge along‐
strike, and from the Late Tertiary to the present day Differences or similarities in the isotopic compo‐
sition of Late Tertiary and Quaternary arc mag‐
mas in Ecuador can be attributed to the influence of crustal basement assimilation if the arc magma source, i.e., the mantle wedge, stayed isotopically 106 homogeneous at this time scale. The latter seems a plausible assumption if related discussions in the Central Andes are taken into account. Sub‐
duction erosion is proposed to have influenced the Sr, Nd, and Pb isotopic composition of the Central Andean mantle wedge implying a chang‐
ing source composition through time (Stern & Skewes 2005). However, these isotopic changes in source composition are shown to be mostly negligible compared to crustal contamination effects, such that the isotopic composition of the mantle wedge can be assumed as broadly con‐
stant during Tertiary‐Quaternary times (Mamani et al. 2010). While the Ecuadorian margin has been partly erosive since the Late Miocene collision with the Carnegie Ridge seamount chain (e.g., Sage et al. 2006), the relative stability of the Ecuadorian arc position during the Tertiary excludes significant earlier subduction erosion (Chapter 2). Poten‐
tially contaminating effects of subduction erosion on the Sr, Nd, and Pb isotopic composition of the mantle wedge are thus expected to be of a much lower magnitude than in parts of the Central An‐
des (Stern & Skewes 2005). This is confirmed by studies of Eocene Macuchi arc magmatism (Chiaradia 2009) and Quaternary arc volcanoes (e.g., Chiaradia et al. 2009) which show that the mantle wedge below Ecuador represents an iso‐
topically broadly homogeneous reservoir, slightly enriched by a sedimentary component, and has not changed systematically in its isotopic compo‐
sition throughout the Tertiary. Latitudinal mantle wedge Pb isotopic heterogeneity along the South American margin only applies at a larger scale, and Tertiary Ecuadorian arc magmas define a sin‐
gle regression line in uranogenic and thorogenic Pb diagrams (Chiaradia & Fontboté 2002). This notion is supported by overlapping isotopic ranges for specific groups of NVZ volcanoes in northern Ecuador and Late Tertiary granitoids in southern and central Ecuador as described above (Fig. 2). We therefore argue that similarities or differ‐
ences in the isotopic composition of Late Tertiary arc magmas in the Western Cordillera and the southern Ecuadorian Sierra, and Quaternary arc volcanoes in northern‐central Ecuador are mainly caused by assimilation of specific crustal base‐
ment lithologies, additionally influenced by vari‐
able crustal thickness. Second‐order modulations of isotopic ratios originating from variations in source (mantle wedge) composition cannot be ruled out, but are not considered as significant at the isotopic scale relevant for the following dis‐
cussion. The role of crustal thickness on iso‐
topic compositions of Tertiary‐
Quarter‐nary Ecuadorian arc magmas Assimilation of oceanic plateau units may drive Tertiary‐Quaternary arc magmas in Ecuador to‐
wards more primitive Sr and Nd isotopic compo‐
sitions than their mantle wedge‐derived parental melts (Chiaradia 2009); assimilation of (partly) continental Eastern Cordillera basement may have the reverse effect, producing complex iso‐
topic patterns of crust‐magma interaction (Chia‐
radia et al. 2009). A thick crust maximizes the likelihood of crustal contamination of arc mag‐
mas (e.g., Hildreth & Moorbath 1988; Annen et al. 2006). However, in a vertically heterogeneous crustal column bulk crustal thickness does not directly scale with a specific contamination signal. Rather, the relative thickness of crustal material of contrasting isotopic composition (here: oce‐
anic plateau vs. Eastern Cordillera basement) and tectonomagmatic controls on the depth of crustal magma evolution in a given crustal column con‐
stitute the dominant control factors for the final isotopic composition of Tertiary‐Quaternary Ec‐
uadorian arc magmas. At what crustal level did Late Tertiary and Quaternary arc magmas acquire their isotopic characteristics? Bulk crustal contamination of evolving arc mag‐
mas by assimilation of crustal lithologies or mix‐
ing with crustal partial melts principally occurs in hot zones at lower to mid‐crustal levels; shallow crustal magma evolution does not involve signifi‐
cant compositional modification of arc magmas by crustal contamination (or, more general, as‐
similation and fractional crystallization; AFC) unless large, supra‐solidus magmatic systems form which need to be sustained by high magma supply rates (Annen et al. 2006). Potential shal‐
low crustal assimilants may be of significantly more radiogenic isotopic compositions than deep to mid‐crustal lithologies, and thus may leave a distinct isotopic fingerprint on arc magmas (Hil‐
dreth & Moorbath 1988; Dungan & Davidson 2004). Consequently, discriminating between the upper and lower crustal contributions to whole‐
rock isotopic compositions of arc magmas is of major importance to delineate the deep through shallow crustal basement architecture. Petrologic studies (e.g., Chiaradia et al. 2009) demonstrate that Quaternary NVZ magmas in northern Ecuador mostly evolved in the stability fields of garnet and amphibole, and outside the stability field of plagioclase, suggesting that these magmas dominantly acquired their crustal iso‐
topic signatures through polybaric evolution at lower to mid‐crustal levels, although minor peri‐
ods of subsequent shallow crustal magma evolu‐
tion do occur for some volcanic centers. Late Ter‐
tiary arc magmas evolved in an overall thinner crust than present‐day arc magmas (Jaillard et al. 2005) and form two distinct groups (Chapter 5): the dominant group comprises most granitoids of the Western Cordillera (Apuela, Junin, Balsa‐
pamba, Telimbela, Gaby, Chaucha), as well as volcanic formations (Saraguro Group at Cañicapa and Tres Chorreras; Quimsacocha) in the south‐
ern‐central Ecuadorian Sierra. REE patterns of this group commonly lack negative Eu anomalies indicating that the parental magmas of these in‐
trusions and volcanics did not fractionate signifi‐
cant amounts of plagioclase at shallow crustal levels (e.g., at <0.4 GPa, corresponding to the maximum pressure where plagioclase precedes amphibole on the liquidus for water‐rich basaltic‐
andesitic melts; Grove et al. 2003). Rather, paren‐
tal melts to these intrusions and volcanics seem to have evolved at deep to mid‐crustal levels without major compositional overprinting by shallow crustal magma evolution (Chapter 5). A second group comprises some intrusions of the Cangrejos‐Zaruma intrusive belt, the bulk of the Saraguro Group (here: at Portovelo and Chaucha), as well as some minor intrusions at El Mozo. REE patterns of this group are usually characterized by minor‐moderate negative Eu anomalies suggesting that these intrusions and volcanics derive from parental magmas which underwent significant plagioclase fractionation at 107
shallow crustal levels (Chapter 5). Late Tertiary arc magmas thus variably acquired their crustal isotopic signatures at deep, mid‐, and shallow crustal levels. Shallow vs. deep to mid‐crustal magma evolution can be qualitatively discriminated using the Sr/Y ratio (e.g., Bachmann et al. 2005); in our dataset, Sr/Y ratios >30 indicate the absence of pro‐
nounced shallow crustal magma evolution in Late Tertiary Ecuadorian arc magmas (Chapter 5). It is important to note, however, that the Sr/Y ratio does not directly scale with the absolute depth of magma evolution, as it can also be affected by pressure‐insensitive accessory phase fractiona‐
tion (Chapter 5). Figure 3 shows the Sr, Nd, and Pb isotopic compositions of NVZ volcanic centers and Late Tertiary samples as a function of Sr/Y ratios, and Sr isotopic compositions as a function of SiO2 (SiO2, Sr, and Y concentrations for Late Tertiary samples from Chapter 5). Individual groups of Quaternary arc volcanoes identified in the previous section mostly define subparallel Figure 3: Diagrams of 87Sr/86Sr, εNdinitial, 206Pb/204Pb, and 207Pb/204Pb vs. Sr/Y, and 87Sr/86Sr vs. SiO2. The Sr/Y ratio serves as a proxy for shallow crustal vs. mid‐ to deep crustal (>30) magma evolution. Arc magmas color‐coded ac‐
cording to the isotopic classification scheme of Figs. 1 and 2, except for southern Ecuadorian Sierra units which are uniformly shown in black. The various groups define broad subparallel isotopic arrays at Sr/Y >30 suggesting that specific groups of arc magmas undergo AFC (assimilation and fractional crystallization) ± mixing processes at deep to mid‐crustal levels involving variable proportions of different basement units. In addition, Late Tertiary arc magmas of the Cangrejos‐Zaruma intrusive belt undergo significant shallow crustal AFC processes and assimilate crustal ma‐
terial characterized by more radiogenic Sr and less radiogenic Nd isotopic compositions than deep to mid‐crustal basement units. Distinct crustal AFC trends exist in a 87Sr/86Sr vs. SiO2 plot, in agreement with the notions above. Sr/Y and SiO2 data from Chapter 5 (Tertiary magmas), and references given in Fig. 1 (NVZ). See text for discussion.
108 isotopic arrays for Sr/Y ratios >30 (the distinction between groups 1 and 2 is somewhat ambiguous: while mostly overlapping in 87Sr/86Sr and εNdinitial, they systematically differ in 207Pb/204Pb vs. Sr/Y plots, but the latter difference is close to the ana‐
lytical resolution). Parental melts to these arc volcanics thus variably assimilated deep to mid‐
crustal basement lithologies of distinct isotopic compositions. Magma differentiation by assimila‐
tion of different crustal lithologies is further demonstrated by an 87Sr/86Sr vs. SiO2 plot show‐
ing distinct crustal AFC trends for various isotopic groups (Fig. 3). As noted above, Quaternary arc volcanoes merge at a common isotopic composition (Fig. 2) corre‐
sponding to the most primitive magma composi‐
tions. This is not directly visible from Figure 3 be‐
cause the Sr/Y ratio of mantle‐derived primitive melts may show a higher variability both across and along the arc, e.g., due to variable amounts of mantle wedge fluxing by slab‐derived fluids, and different degrees of partial melting (Chiara‐
dia et al. 2009). Western Cordillera granitoids define a further isotopic subgroup at Sr/Y ratios >30; as noted above, these magmas acquire their crustal iso‐
topic signatures by assimilation of primitive CCOP units at depth. Late Tertiary intrusions and vol‐
canics of southern Ecuador overlap with the iso‐
topic range defined by arc volcanoes east of the PF (group 3) at Sr/Y ratios >30, but show more radiogenic Sr and less radiogenic Nd ratios at Sr/Y <30 (Fig. 3). Consequently, we infer that the iso‐
topic characteristics of Late Tertiary magmas in southern Ecuador partly reflect deep to mid‐
crustal assimilation of basement units similar to arc volcanoes east of the PF in northern Ecuador, but they may additionally acquire distinct isotopic signatures by further shallow crustal magma evo‐
lution. Note that the isotopic composition of the El Mozo intrusions is rather primitive (Fig. 2) sug‐
gesting that, despite showing minor negative Eu anomalies, isotopic contamination of El Mozo parental magmas by upper crust material was limited such that their lower to mid‐crustal iso‐
topic signatures were preserved. Tectonic implications of Sr‐Nd‐Pb iso‐
topic systematics in Late Tertiary and Quaternary arc magmas The isotopic similarities between Quaternary northern Ecuadorian volcanic centers east of the PF and Late Tertiary southern Ecuadorian intru‐
sions and volcanics at Sr/Y >30 (Figs. 1‐3) suggest that the isotopic compositions of lower to mid‐
crustal IAD basement units north and east of the Jubones fault (Fig. 1) are broadly homogeneous in along‐arc dimension. In contrast, isotopic data for the Cangrejos‐Zaruma intrusive belt record both shallow and deep to mid‐crustal assimilation of continental crust‐dominated basement litholo‐
gies (high radiogenic Sr and low radiogenic Nd); the latter data support the notion that the north‐
ern limit of deep crustal basement units of the El Oro micro‐continental block is bracketed be‐
tween the Piñas‐Portovelo and Jubones faults as significantly more primitive isotopic compositions occur only north of the Jubones fault. Isotopic compositions show a systematic across‐
arc distribution in northern Ecuador. In the frame of a basement architecture characterized by jux‐
taposed oceanic and continental basement do‐
mains (as applicable for Ecuador; e.g., Jaillard et al. 2005) of strongly contrasting isotopic compo‐
sitions (compare reference fields in Fig. 2), more radiogenic Sr and less radiogenic Nd isotopic compositions of arc magmas are mainly indicative of increasing continental signatures imposed on evolving magmas by a higher proportion of conti‐
nental versus oceanic basement assimilation in a given crustal column. The role of Pb isotopes is more difficult to define as oceanic plateau mate‐
rial in Ecuador is characterized by a wide range in 206
Pb/204Pb (Fig. 2) such that the latter isotopic ratio cannot discriminate oceanic vs. continental material on a regional scale; more radiogenic 207
Pb/204Pb compositions, on the other hand, are indicative of a stronger continental basement signature, but isotopic variations are of such a small scale that they approach the analytical resolution limit (Fig. 2). Figure 4 illustrates a schematic cross section of the Ecuadorian arc at c. 0.5°S based on seismic studies (Guillier et al. 2001; Jaillard et al. 2005). Eastward underthrusting of high‐density oceanic plateau material below the IAD results in east‐
109
ward thickening of the proportion of Eastern Cordillera basement (simplified here as a single unit) relative to oceanic plateau material in a given crustal column. Consequently, the potential role of Cordillera Real basement as an assimilant for NVZ magmas evolving in lower to mid‐crustal hot zones progressively increases eastwards. This regional underthrusting of oceanic plateau mate‐
rial below the IAD is mirrored by across‐arc trends in Sr, Nd, and Pb isotopic compositions of NVZ arc volcanoes which get progressively more continental in character towards the east (Figs. 1, 4). Initial magma differentiation in a hot zone at the base of the crust (or in the uppermost litho‐
spheric mantle) would mostly include oceanic plateau material, whereas subsequent mid‐
crustal magma processing should progressively involve increasing amounts of Eastern Cordillera basement units. In contrast, parental magmas to Western Cordillera granitoids were entirely con‐
fined to oceanic plateau material (or arc intrusive root zones) during their crustal transit such that arc magmas evolve towards less radiogenic Sr and more radiogenic Nd isotopic compositions (compare Fig. 2). Tertiary and NVZ isotopic com‐
positions in northern Ecuador, therefore, are consistent with the crustal structure inferred from seismic studies involving regional under‐
thrusting of oceanic plateau material below the paleo‐continental margin (Guillier et al. 2001). In addition to considering the differing isotopic characteristics of main arc volcanoes east and west of the Peltetec fault zone purely as a func‐
tion of a transitional change of deep to mid‐
crustal, continental versus oceanic basement units, the presence of different continental (Li‐
therland et al. 1994) or oceanic terrane units (such as a second, pre‐Late Cretaceous oceanic plateau fragment; e.g., Mamberti et al. 2003) might further influence the isotopic signature of arc magmas produced at these volcanoes, al‐
though the geometry of their deep to mid‐crustal extensions cannot be predicted with the current dataset. Ecuadorian frontal arc volcanoes such as Pichin‐
cha are characterized by the presence of crustal xenoliths whose high‐grade (granulite) metamor‐
phic character implies a mid‐ to lower crustal ori‐
gin (Chiaradia et al. 2009). Most xenoliths are characterized by low radiogenic Sr and high ra‐
110 diogenic Nd isotopic compositions overlapping with basement units cropping out in the Western Cordillera (amphibolites of Amortegui 2007; Fig. 2). A single Pichincha xenolith detected by Chiaradia et al. (2009) has an isotopic composi‐
tion resembling shallow crustal metapelites of the Western Cordillera (“Western Cordillera shal‐
low basement” in Fig. 2; Amortegui 2007). Incor‐
poration of this type of xenolith into Pichincha melts might thus have occurred at shallow crustal levels. Alternatively, if this xenolith type repre‐
sents a deep to mid‐crustal lithology, its occur‐
rence might be reconciled with the generalized model presented in Figure 4 if non‐vertical magma ascent along transcrustal fault systems such as the CPPF is taken into account. Structur‐
ally controlled melt ascent including significant non‐vertical components has been documented elsewhere, e.g., by pluton emplacement along fault ramps in the Sevier fold‐and‐thrust belt (Ka‐
lakay et al. 2001). High‐resolution seismic data imaging crustal structures are not available in southern Ecuador, but a fundamental change in basement architec‐
ture is implied by a major change in structural trends from N‐S (central‐northern Ecuador) to E‐
W (El Oro massif south of the Jubones fault in southern Ecuador; Fig. 1), and is further clearly visible in the isotopic compositions of granitoids of the Cangrejos‐Zaruma intrusive belt (see above). In this area, the western surface trace of the oceanic plateau‐bounding suture zone (the CPPF) splays off and intersects the Western Cor‐
dillera towards the Gulf of Guayaquil. Instead of displaying a transition in isotopic compositions across the arc as in northern Ecuador, there seems to be a rather sharp change in Sr, Nd, and Pb isotopic compositions from the west (Gaby intrusions) to the east (Quimsacocha volcano, Chaucha intrusions, Saraguro volcanics) of the CPPF. This might suggest that, concomitant with its change in strike direction, the mid‐ to deep crustal structure of the suture zone possibly changes from 35° E‐dipping in northern‐central Ecuador towards a more subvertical orientation in southern Ecuador. An alternative interpreta‐
tion of the relatively sharp isotopic contrast across the CPPF in southern Ecuador might be the absence of an additional terrane unit equivalent to the NVZ basement west of the Peltetec fault zone as discussed above. Based on gravimetric data, Feininger & Seguin (1983) modeled crustal thickness profiles in Ec‐
uador using a strictly vertical dip for the CPPF (their “Romeral” fault), and confined oceanic pla‐
teau material to the west of the CPPF. Conse‐
quently, these authors identified the absence of the IAD in southern Ecuador (with its low‐density volcaniclastic infill) as a major influence on the gravimetric anomaly pattern, and suggested that the crustal thickness in southern Ecuador might exceed the thickness in northern Ecuador. When taking into account a more realistic 35° E dip of the CPPF in northern Ecuador implying regional underthrusting of high‐density oceanic plateau material, highly negative gravity anomalies (down to ‐292 mgal) in the Andean region east of the CPPF might necessitate an even thicker crustal root in northern Ecuador, making it unlikely that the crustal thickness of southern Ecuador gener‐
ally exceeds northern Ecuador (50‐70 km; Guillier et al. 2001). Rather, the crustal thickness of the main Andean root in southern Ecuador might be similar to or slightly thinner than in northern Ec‐
uador, in agreement with data from northern Peru where a maximum crustal thickness of 45 km below the Western Cordillera is inferred (Fu‐
kao et al. 1989). Figure 4: Schematic crustal cross section of northern Ecuador, and 87Sr/86Sr, εNdinitial, 206Pb/204Pb, and
207
Pb/204Pb across‐arc distribution trends for NVZ vol‐
canoes, the Apuela batholith and the Junin porphyry intrusions between 1°S to 0.5°N, orthogonally pro‐
jected onto the cross section shown in Fig. 1; same
symbol key as in Fig. 3. Crustal section simplified and modified from Jaillard et al. (2005) sketching under‐
thrusting of oceanic plateau units (v pattern) below the
paleo‐continental margin (gray). For simplicity, Eastern
Cordillera basement is visualized here as a single unit,
but in detail consists of multiple meta‐sedimentary and meta‐igneous units possibly representing different ter‐
ranes of both oceanic and continental affinity (Litherland et al. 1994). Orange bars indicate deep to mid‐crustal hot zones where major magma evolution occurs (Fig. 3; Annen et al. 2006), progressively including a higher component of Eastern Cordillera basement towards the east, and most clearly reflected by decreasing εNdinitial, and increasing 207
Pb/204Pb across the arc. Note that transcrustal magma ascent might partly be focused along non‐vertical struc‐
tures such as the Chimbo‐Toachi shear zone (CTSZ) and the Calacali‐Pujili‐Pallatanga fault (CPPF), and, possibly, the Peltetec fault (PF), allowing incorporation of crustal xenoliths of continental crust affinity into frontal arc (Pichin‐
cha) magmas. Section not vertically exaggerated. 111
Conclusions Late Tertiary‐Quaternary arc magmas in Ecuador derive from an isotopically broadly homogeneous mantle wedge, and acquire variable deep, mid‐, and shallow crustal Sr, Nd, and Pb isotopic signa‐
tures during subsequent magma evolution stages. These isotopic imprints provide insights into the crustal basement architecture of Ecua‐
dor’s Western Cordillera, the IAD, and the south‐
ern Ecuadorian Sierra region. The Ecuadorian crust in the investigated areas seems to be vertically heterogeneous where the relative thickness of crustal material of contrast‐
ing isotopic composition (oceanic plateau vs. Eastern Cordillera basement) and tectonomag‐
matic controls on the depth of crustal magma evolution in a given crustal column constitute the dominant control factors for the isotopic compo‐
sition of Ecuadorian arc magmatic products. Eastern Cordillera basement) and tectonomag‐
matic controls on the depth of crustal magma evolution in a given crustal column constitute the dominant control factors for the isotopic compo‐
sition of Ecuadorian arc magmatic products. Crustal imprints on arc magmas are mostly rela‐
tively primitive in Sr and Nd, and highly variable in Pb, and define a regionally systematic distribu‐
tion pattern, consistent with the notion of large‐
scale underthrusting of allochthonous oceanic plateau material below the paleo‐continental margin of northern‐central Ecuador as inferred from seismic studies (Guillier et al. 2001). With increasing distance from the trench, Late Terti‐
ary‐Quaternary arc magmas evolve towards pro‐
gressively more radiogenic 87Sr/86Sr and 207
Pb/204Pb, and less radiogenic 143Nd/144Nd com‐
positions at deep to mid‐crustal levels. Late Tertiary arc magmas of the southern Ecua‐
dorian Sierra east of the CPPF overlap with iso‐
topic compositions of Quaternary arc volcanoes east of the PF in northern Ecuador suggesting along‐strike continuity of similar deep to mid‐
crustal basement units. Similarly, frontal arc vol‐
canoes in northern Ecuador isotopically overlap with the composition of the Gaby intrusive center in southern Ecuador west of the Bulubulu fault system (an eastern splay fault of the CPPF). Granitoids of the Cangrejos‐Zaruma intrusive belt in southern Ecuador are characterized by highly radiogenic Sr and Pb, and low radiogenic Nd iso‐
topic compositions, which have no equivalent in central‐northern Ecuadorian arc magmas. Their isotopic compositions relate to the mostly shal‐
low crustal assimilation of basement units form‐
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114 Winkler, W., Villagomez, D., Spikings, R., Abegglen, P, Tobler, S, Eguez, A. (2005): The Chota basin and its significance for the inception and tectonic setting if the inter‐Andean depression in Ecuador. Journal of South American Earth Sciences 19; 5‐19. Wörner, G., Moorbath, S., Harmon, R. S. (1992): An‐
dean Cenozoic volcanic centers reflect basement iso‐
topic domains. Geology 20; 1003‐1006. Appendix: Data tables Table A1: Whole rock 87Sr/86Sr, 143Nd/144Nd, 206Pb/204Pb, 207Pb/204Pb, and 208Pb/204Pb compositions of North‐
ern Volcanic Zone volcanoes 115
87
86
143
Table A1: Whole rock Sr/ Sr, Nd/
Northern Volcanic Zone volcanoes
87
Sample
Magmatic
Sr
center
86
Sr
144
Nd,
206
Pb/
143
Nd
204
Pb,
εNd
144
Nd
207
Pb/
204
208
Pb, and
204
Pb/
Pb compositions of
206
207
208
latitude
204
204
204
(projected)
Pb
Pb
Pb
Pb
Pb
Pb
ref*
Frontal arc
E05001
Pululagua
0.70420
0.51286
4.4
18.889
15.587
38.552
78.60
1
E05003
Pululagua
0.70416
0.51289
4.8
18.868
15.596
38.565
78.60
1
E05007
Pululagua
0.70417
0.51290
5.0
18.906
15.604
38.619
78.60
1
E05008
Pululagua
0.70415
0.51293
5.7
18.912
15.587
38.566
78.60
1
E05042
Pululagua
0.70412
0.51291
5.4
18.933
15.609
38.645
78.60
1
Pul-9
Pululagua
0.70412
0.51291
5.4
18.934
15.605
38.610
78.60
2
Pul-11
Pululagua
0.70414
0.51292
5.5
18.893
15.606
38.562
78.60
2
Pul-4
Pululagua
0.70414
0.51295
6.2
18.921
15.595
38.572
78.60
2
Pul-7
Pululagua
0.70414
0.51289
5.1
18.921
15.596
38.572
78.60
2
E05016
Pichincha
0.70408
0.51287
4.6
18.957
15.575
38.690
78.60
1
E05017
Pichincha
0.70399
0.51286
4.4
19.018
15.589
38.728
78.60
1
E05018
Pichincha
0.70410
0.51287
4.5
18.884
15.596
38.578
78.60
1
E05130
Pichincha
0.70409
0.51293
5.8
18.927
15.596
38.638
78.60
1
E05131
Pichincha
0.70407
0.51286
4.3
18.934
15.597
38.631
78.60
1
E05010
Pichincha
0.70404
0.51290
5.2
19.017
15.593
38.780
78.60
1
E05012
Pichincha
0.70400
0.51289
4.9
18.957
15.575
38.690
78.60
1
E05013
Pichincha
0.70401
0.51287
4.5
18.999
15.577
38.738
78.60
1
E05014
Pichincha
0.70404
0.51292
5.5
18.945
15.589
38.654
78.60
1
E05015
Pichincha
0.70410
0.51289
4.9
18.933
15.582
38.630
78.60
1
PICH 4C
Pichincha
0.70404
0.51288
4.7
19.001
15.591
38.723
78.60
3
PICH 9C
Pichincha
0.70407
0.51288
4.8
18.984
15.577
38.670
78.60
3
PICH 10
Pichincha
0.70406
0.51289
5.0
19.000
15.584
38.708
78.60
3
Gp-1
Pichincha
0.70395
0.51289
4.9
18.958
15.597
38.698
78.60
2
Ql-12
Quilotoa
0.70403
0.51286
4.4
18.985
15.644
38.750
78.65
2
ATAC 2C
Atacazo
0.70420
0.51289
5.0
18.878
15.589
38.521
78.55
3
ATAC 8
Atacazo
0.70422
0.51286
4.4
18.962
15.601
38.627
78.55
3
ATAC 12B
Atacazo
0.70434
0.51286
4.4
18.967
15.600
38.604
78.55
3
AT01
Atacazo
0.70419
0.51285
4.1
n/a
n/a
n/a
78.55
4
8557 At
Atacazo
0.70430
0.51289
4.9
n/a
n/a
n/a
78.55
4
At-02
Atacazo
0.70408
0.51289
5.0
18.933
15.602
38.619
78.55
2
957 g
Atacazo
0.70430
0.51288
4.7
n/a
n/a
n/a
78.55
2
Frontal arc basement xenoliths
E05008a
Pululagua
0.70415
0.51289
4.9
18.862
15.552
38.473
78.60
1
E05130a
Pichincha
0.70392
0.51288
4.8
18.861
15.588
38.587
78.60
1
E05011
Pichincha
0.70686
0.51260
-0.7
18.607
15.651
38.548
78.60
1
E05015a
Pichincha
0.70401
0.51290
5.1
18.945
15.611
38.699
78.60
1
Main arc (west of Peltetec fault)
Imb-1
Imbabura
0.70408
0.51290
5.2
19.098
15.600
38.770
78.40
2
Imb-3
Imbabura
0.70408
0.51291
5.3
19.099
15.606
38.786
78.40
2
Imb-11
Imbabura
0.70407
0.51293
5.7
19.072
15.611
38.766
78.40
2
Imb-26
Imbabura
0.70412
0.51288
4.7
18.921
15.607
38.614
78.40
2
Imb-28
Imbabura
0.70406
0.51292
5.5
18.953
15.616
38.657
78.40
2
Imb-29
Imbabura
0.70411
0.51289
5.0
18.974
15.629
38.717
78.40
2
Imb-30
Imbabura
0.70420
0.51282
3.7
18.913
15.622
38.677
78.40
2
Imb-36
Imbabura
0.70402
0.51291
5.3
19.123
15.617
38.833
78.40
2
116 Table A1 (continued)
Sample
Magmatic
87
143
center
86
144
Sr
Sr
Nd
εNd
Nd
206
207
208
latitude
204
204
204
(projected)
Pb
Pb
Pb
Pb
Pb
Pb
ref*
Imb-38
Imbabura
0.70425
0.51282
3.6
18.900
15.630
38.712
78.40
2
Imb-45
Imbabura
0.70402
0.51290
5.1
19.073
15.618
38.784
78.40
2
E05132
Ilalo
0.70414
0.51282
3.5
18.951
15.612
38.695
78.40
1
E05133
Ilalo
0.70409
0.51296
6.2
18.955
15.621
38.717
78.40
1
E05134
Ilalo
0.70405
0.51285
4.0
18.940
15.606
38.647
78.40
1
E05046
Ilalo
0.70406
0.51284
3.9
18.940
15.607
38.652
78.40
1
Il-2
Ilalo
0.70406
0.51286
4.4
n/a
n/a
n/a
78.40
2
Cx-9
Cotopaxi
0.70414
0.51282
3.7
18.963
15.631
38.710
78.30
2
Cx-10
Cotopaxi
0.70419
0.51284
4.0
18.951
15.637
38.716
78.30
2
5
Main arc (east of Peltetec fault)
CAY56
Cayambe
0.70431
0.51274
2.0
n/a
n/a
n/a
78.20
CAY55B
Cayambe
0.70441
0.51274
2.1
n/a
n/a
n/a
78.20
5
CAY106B
Cayambe
0.70442
0.51264
0.0
n/a
n/a
n/a
78.20
5
CAY107
Cayambe
0.70442
0.51272
1.6
n/a
n/a
n/a
78.20
5
CAY31
Cayambe
0.70446
0.51273
1.7
n/a
n/a
n/a
78.20
5
CAY80A
Cayambe
0.70451
0.51271
1.3
n/a
n/a
n/a
78.20
5
CAY78A
Cayambe
0.70454
0.51262
-0.4
n/a
n/a
n/a
78.20
5
CAY168D
Cayambe
0.70434
0.51280
3.1
n/a
n/a
n/a
78.20
5
CAY98
Cayambe
0.70432
0.51280
3.2
n/a
n/a
n/a
78.20
5
CAY8
Cayambe
0.70432
0.51277
2.5
n/a
n/a
n/a
78.20
5
CAY179A
Cayambe
0.70445
0.51268
0.9
n/a
n/a
n/a
78.20
5
CAY39
Cayambe
0.70439
0.51276
2.3
n/a
n/a
n/a
78.20
5
CAY87
Cayambe
0.70443
0.51274
2.1
n/a
n/a
n/a
78.20
5
CAY44A
Cayambe
0.70437
0.51281
3.3
n/a
n/a
n/a
78.20
5
CAY46B
Cayambe
0.70442
0.51277
2.6
n/a
n/a
n/a
78.20
5
CAY45C
Cayambe
0.70408
0.51270
1.3
n/a
n/a
n/a
78.20
5
ANT 54
Antisana
0.70454
0.51273
1.8
18.895
15.637
38.711
78.20
3
ANT 26
Antisana
0.70441
0.51274
2.0
18.886
15.654
38.758
78.20
3
ANT 29C
Antisana
0.70449
2.2
18.974
15.691
38.914
78.20
3
ANT8
Antisana
0.70424
0.51275
0.51276
2.4
n/a
n/a
n/a
78.20
6
ANT10
Antisana
0.70440
0.51273
1.8
n/a
n/a
n/a
78.20
6
ANT14C
Antisana
0.70440
0.51274
1.9
n/a
n/a
n/a
78.20
6
ANT28
Antisana
0.70439
0.51275
2.1
n/a
n/a
n/a
78.20
6
ANT32
Antisana
0.70445
0.51273
1.7
n/a
n/a
n/a
78.20
6
ANT36
Antisana
0.70431
0.51273
1.9
n/a
n/a
n/a
78.20
6
ANT37
Antisana
0.70432
0.51275
2.2
n/a
n/a
n/a
78.20
6
ANT46
Antisana
0.70438
0.51273
1.8
n/a
n/a
n/a
78.20
6
ANT47
Antisana
0.70437
0.51274
1.9
n/a
n/a
n/a
78.20
6
ANT60
Antisana
0.70434
0.51273
1.7
n/a
n/a
n/a
78.20
6
ANT61
Antisana
0.70437
0.51272
1.6
n/a
n/a
n/a
78.20
6
ANT62
Antisana
0.70426
0.51276
2.4
n/a
n/a
n/a
78.20
6
3.2An
Antisana
0.70421
0.51280
3.1
n/a
n/a
n/a
78.20
4
HHJ-An
Antisana
0.70419
0.51280
3.1
n/a
n/a
n/a
78.20
4
GS-3
Antisana
0.70420
0.51289
5.0
n/a
n/a
n/a
78.20
4
3D2
Antisana
0.70424
0.51289
5.0
n/a
n/a
n/a
78.20
4
3D1 An
Antisana
0.70414
0.51285
4.1
18.936
15.618
38.728
78.20
2
2G3T-A
Antisana
0.70437
0.51275
2.3
18.924
15.643
38.762
78.20
2
HHV
Antisana
0.70416
0.51279
3.1
18.985
15.703
38.978
78.20
2
117
Table A1 (continued)
Sample
Magmatic
87
143
center
86
144
Sr
Sr
Nd
εNd
Nd
206
207
208
latitude
204
204
204
(projected)
Pb
Pb
Pb
Pb
Pb
Pb
ref*
E05035
Chacana
0.70434
0.51273
1.9
18.979
15.617
38.705
78.15
1
E05137
Chacana
0.70428
0.51277
2.6
18.933
15.625
38.704
78.15
1
E05138
Chacana
0.70427
0.51277
2.6
18.920
15.641
38.762
78.15
1
E05142
Chacana
0.70430
0.51278
2.7
18.880
15.612
38.640
78.15
1
E05143
Chacana
0.70423
0.51276
2.4
18.901
15.633
38.722
78.15
1
E05144
Chacana
0.70423
0.51279
2.9
18.889
15.604
38.631
78.15
1
E05147
Chacana
0.70426
0.51273
1.8
18.912
15.610
38.675
78.15
1
E05141
Chacana
0.70445
0.51270
1.3
18.953
15.602
38.671
78.15
1
E05030
Chacana
0.70409
0.51281
3.3
18.977
15.632
38.781
78.15
1
E05036
Chacana
0.70414
0.51280
3.2
18.911
15.626
38.723
78.15
1
E05135
Chacana
0.70426
0.51275
2.2
18.938
15.639
38.768
78.15
1
E05136
Chacana
0.70419
0.51273
1.8
18.927
15.638
38.769
78.15
1
E05019
Chacana
0.70430
0.51276
2.4
18.897
15.624
38.688
78.15
1
E05021
Chacana
0.70430
0.51274
2.0
18.953
15.646
38.799
78.15
1
E05032
Chacana
0.70427
0.51276
2.5
18.916
15.634
38.743
78.15
1
E05025
Chacana
0.70464
0.51255
-1.8
19.057
15.654
38.872
78.15
1
E05028
Chacana
0.70456
0.51275
2.2
18.933
15.624
38.748
78.15
1
E05139
Chacana
0.70465
0.51269
1.1
19.008
15.650
38.833
78.15
1
E05140
Chacana
0.70465
0.51271
1.4
19.026
15.626
38.763
78.15
1
Ch-3
Chacana
0.70404
0.51284
4.0
18.956
15.613
38.680
78.15
2
Ch-6
Chacana
0.70432
0.51277
2.6
18.983
15.633
38.755
78.15
2
Ch-7
Chacana
0.70423
0.51276
2.5
19.028
15.688
38.947
78.15
2
Ch-8
Chacana
0.70427
0.51280
3.3
18.925
15.646
38.694
78.15
2
Ch-4
Chacana
n/a
0.51271
1.5
18.925
15.646
38.775
78.15
2
Ch-5
Chacana
0.70434
0.51278
2.7
18.918
15.632
38.732
78.15
2
Sg-13
Sangay
0.70437
0.51274
2.0
18.816
15.644
38.754
n/a
2
Cl-11
Chalupas
0.70443 0.51274
2.0
18.993
15.668 38.873
78.10
*References: 1 - Chiaradia et al. (2009); 2 - Bryant et al. (2006); 3 - Bourdon et al. (2003); 4 - Barragan et al.
(1998); 5 - Samaniego et al. (2005); 6 - Bourdon et al. (2002)
118 2
CHAPTER V ADAKITE‐LIKE FEATURES IN LATE OLIGOCENE TO LATE MIOCENE EC‐
UADORIAN ARC MAGMAS AND THEIR SIGNIFICANCE FOR PORPHYRY‐
RELATED ORE DEPOSITS Abstract This study presents a comprehensive dataset of the geochemical composition of Late Oligocene to Late Miocene intrusions associated in space and time with porphyry‐related ore deposits in Ecuador, supple‐
mented by compositional data on several arc volcanic formations of the same age. Our aim is to describe the spatio‐temporal distribution pattern of adakite‐like geochemical features related to Late Tertiary arc magmatism, and explore its significance for porphyry‐related mineralization in Ecuador. Most investigated intrusions are moderately to highly differentiated hornblende‐ ±biotite‐bearing tonalites, granodiorites, and quartz‐diorites, and often form part of larger Oligocene‐Miocene batholith complexes; arc volcanics represent mostly flows and subordinate tuffs of andesitic and dacitic‐rhyolitic composition. The overall spatio‐temporal distribution of adakite‐like features in Ecuadorian arc magmas is semi‐
systematic; the relative proportion of adakite‐like magmas increases with decreasing age, and is higher in northern‐central than in southern Ecuador. Magmatic centers characterized by (partly) adakite‐like magma‐
tism are mainly hosted by the Western Cordillera and comprise Balsapamba (c. 21 Ma), Apuela‐Junin (13‐6 Ma), Chaucha (c. 10 Ma), and Quimsacocha (7 Ma). High Sr/Y ratios (the commonly used main criteria to signal adakite‐like magma compositions) of Late Tertiary Ecuadorian arc magmas are mainly derived from strong Y (along with heavy REE) depletion of their parental melts at broadly constant Sr contents, and are related to fractionation/restite equilibration effects of amphibole, garnet, and titanite. In Early to Mid‐Miocene Ecuadorian arc magmas, amphibole (± accessory titanite) fractionation/restite equilibration caused silicic melts to evolve towards adakite‐like compositions; combined amphibole and garnet fractionation/restite equilibration is only observed in the Late Miocene Quimsacocha volcanic cen‐
ter, and continues to the present day. While Y depletion by amphibole fractionation/restite equilibration is particularly efficient for silicic melt compositions, garnet fractionation/restite equilibration produces strong Y depletion already in more mafic melt compositions, i.e, during earlier differentiation stages. Significant shallow crustal plagioclase fractionation affects some, but not all Late Tertiary arc magmas in southern Ec‐
uador; it is of minor petrogenetic significance for Miocene intrusions of the Western Cordillera in northern‐
central Ecuador. A preferential association of adakite‐like features with a specific basement lithology can‐
not be observed. Systematic trace element variations (Sr, Y, REE) through time are indicative of progressively increasing high‐
pressure arc magma differentiation from the Late Oligocene to the Late Miocene, either by crustal thicken‐
ing, or/and by the downwards migration of crustal hot zones. However, adakite‐like features are also locally observed in the Early Miocene in a regional petrogenetic setting otherwise dominated by low‐pressure magma evolution. Where porphyry‐related ore deposits are associated with batholith complexes recording multi‐m.y. precur‐
sor magmatism, porphyry emplacement commonly represents a late intrusive event; in that case porphyry parental melts tend to evolve towards more adakite‐like compositions than precursor batholith intrusions indicating downward migration of the focus of crustal magma evolution towards greater depth and/or in‐
creasing water contents in the magmatic system. However, it is important to note that these compositional changes between porphyry and host intrusions mostly reflect broad changes in arc magma composition through time at a regional scale. Systematic compositional changes between porphyry and precursor intru‐
119
sions are not recorded if their relative emplacement age difference is small. The fact that porphyry‐related ore deposits in Ecuador formed throughout the Late Oligocene to Late Miocene (24‐6 Ma) over a large lati‐
tudinal range (c. 0° to 3°30’S) supports the notion that any arc magma of a sufficient volume has the poten‐
tial to form porphyry‐related mineralization. In some cases adakite‐like magmatism may, however, reflect favorable tectonomagmatic preconditioning of porphyry parental melts for subsequent porphyry‐related mineralization. Introduction Active arc volcanism in the Andean chain, where the Nazca plate is subducting below the western plate edge of South America, clusters in three major zones referred to as Northern, Central, and Southern Volcanic Zone (NVZ, CZV, and SVZ, re‐
spectively). These zones are separated by arc segments extending along‐arc for several 100 km which are characterized by the absence of active arc volcanism (Fig. 1). Voluminous Tertiary arc magmatic products extend beyond the present‐
day NVZ both in along‐ and across‐arc dimension. While the geochemical features of Pleistocene‐
Holocene NVZ volcanism of the Ecuadorian Andes have been extensively studied (e.g., Bourdon et al. 2003; Garrison & Davidson 2003; Samaniego et al. 2005; Bryant et al. 2006; Garrison et al. 2006; Chiaradia et al. 2009a), geochemical data on Tertiary arc magmatism are sparse. Building on a pilot study by Chiaradia et al. (2004), this contribution presents new geochemical data for a number of Late Tertiary intrusions and volcanic formations of the northern‐central Ecuadorian Western Cordillera and the Interandean Sierra region of southern Ecuador, representing the most extensive dataset to date for Tertiary igne‐
ous rocks in the northern Andes. We focus on investigating the occurrence of adakite‐like fea‐
tures in Tertiary arc magmas associated with porphyry‐related ore deposits, and evaluate their metallogenetic significance. Tertiary‐Quaternary adakite‐like magmatism in Ecuador A main aspect addressed by most studies on pre‐
sent‐day NVZ volcanism is the mechanism to generate adakite‐like geochemical signatures in Ecuadorian arc volcanics. Adakite‐like magmatism is characterized by a specific geochemical compo‐
sition indicative of, amongst others, parental melt equilibration with a Y‐ (and heavy rare earth element; HREE‐) retentive mineral phase, and absence of significant plagioclase fractionation (see review by Richards & Kerrich 2007, and ref‐
erences therein). The term “adakite‐like” was introduced by Richards & Kerrich (2007) for arcs built on thick continental crust such as the Andes to avoid confusion with adakite sensu stricto which is exclusively defined for island arc settings (Defant & Drummond 1990). In this context, the term "adakite‐like" does not carry any specific petrogenetic implication with respect to the magma source. In particular, adakite‐like chemi‐
cal features of continental arc rocks are not nec‐
essarily associated with slab melting as adakites sensu stricto potentially are (Richards & Kerrich 2007). Generation of adakite‐like geochemical composi‐
tions in present‐day NVZ magmas is envisaged either by slab melting and mantle wedge con‐
tamination by slab melting of the subducting young (<24 Ma), hot Nazca slab (e.g., Gutscher et al. 2000; Bourdon et al. 2003; Samaniego et al. 2005), or high‐pressure equilibration of arc mag‐
mas at deep to mid‐crustal levels (e.g., Garrison Figure 1 (next page): Geological map of Tertiary‐Quaternary Ecuadorian arc units and magmatic centers investigated in this study. Inset shows distribution of the Northern (NVZ), Central (CVZ), and Southern Volcanic Zone (SVZ) resulting
from subduction of the Nazca plate below South America; gray bars indicate positions of major seamount chains pres‐
ently colliding with the central‐southern American margin (from N to S: Cocos, Carnegie, Nazca, and Juan Fernandez
Ridge). Simplified from various references in Chapter 2. 120 121
& Davidson 2003; Garrison et al., 2006; Chiaradia et al. 2009a). Bryant et al. (2006) note the diffi‐
culty to clearly discriminate these two processes, and present intra‐mantle wedge partial melting processes in the garnet stability field as an addi‐
tional option. The Late Miocene collision of the Carnegie Ridge seamount chain with the northern Ecuadorian margin is considered to exert a major influence on arc magmatism either by increasing the geothermal gradient along the subducting slab surface thus facilitating slab melting (Sama‐
niego et al. 2005), or by potentially affecting far‐
field stress and transcrustal magma ascent kinet‐
ics (Chiaradia et al. 2009a). Seismic studies at the Ecuadorian margin demonstrate a continuously subducting Nazca/Farallon slab, subducting at an angle of 25‐30° down to at least 150‐200 km depth (Guillier et al. 2001) such that earlier sug‐
gestions proposing slab melting due to slab flat‐
tening in response to Carnegie Ridge subduction (e.g., Gutscher et al. 2000; Beate et al. 2001) can be excluded. The occurrence of adakite‐like magmatic features in the Ecuadorian subduction system through time is not clearly understood. Beate et al. (2001), Somers et al. (2005), and Amortegui (2007) demonstrate that adakite‐like rock com‐
positions already occur in Late Miocene intru‐
sions and volcanics. Chiaradia et al. (2004) show that a number of Tertiary, pre‐Late Miocene in‐
trusions and volcanic formations lack adakite‐like features, and note an apparent mutual exclusivity of mainly Early to Mid‐Miocene porphyry‐related mineralization and Late Miocene to Holocene adakite‐like magmatism in Ecuador. Recently, Chiaradia (2009) presented evidence for adakite‐
like features of small intrusive bodies of the Eo‐
cene Macuchi island arc sequence of central‐
northern Ecuador. The present study aims to contribute to a better understanding of the spatio‐temporal distribution of adakite‐like magmatism as part of the general petrogenetic evolution of Oligocene‐Miocene Ecuadorian arc magmatism. In particular, we ex‐
plore the metallogenic significance of adakite‐like magmatism for porphyry‐related mineralization in Ecuador, as spatial associations of intrusions with adakite‐like geochemical features and por‐
phyry‐related ore deposits have been observed elsewhere (e.g., Thiéblemont et al. 1997). Conse‐
quently, magmatic centers of this study were se‐
lected based on their spatial association with porphyry‐related ore deposits. The widespread occurrence of this ore deposit type in Ecuador (Prodeminca 2000) ensures a regionally represen‐
tative coverage of Late Tertiary arc magmatism. Parallel geochronologic (Chapters 2, 3) and iso‐
topic (Chapter 4) studies provide age and base‐
ment control, enabling us to calibrate geochemi‐
cal changes in arc magmatism through time and across different crustal basement domains. Regional geology and geody‐
namic setting The Ecuadorian margin is characterized by the typical principal geologic features of a collisional continental arc comprising a foreland basin‐
hosting back‐arc region, a major orogen, split into the Eastern and Western Cordillera, and a forearc sliver which undergoes trench‐parallel, dextral strike‐slip displacement relative to the continent as a result of oblique plate convergence between the Nazca plate and South America (Litherland et al. 1994; Ego et al. 1996). Basement units of the back‐arc region and the Eastern Cordillera are of Precambrian‐Paleozoic age; they are intruded by a voluminous Triassic‐Jurassic arc sequence Figure 2 (next page): Macro‐photographs (A, G) and micro‐photographs (B‐F, H) illustrating mineralogical features rep‐
resentative for analyzed samples of Late Tertiary magmatic centers in Ecuador. A – hornblende‐ and biotite‐bearing granodiorite with weak sericite‐chlorite alteration (Chaucha batholith). B – plagioclase and hornblende‐phyric dacite with fresh glassy matrix (Quimsacocha). C – plagioclase‐hornblende porphyry with resorbed quartz and weak propylitic
alteration (Cangrejos). D – hornblende‐plagioclase porphyry with sodic‐calcic alteration (Gaby). E – hornblende quartz‐
diorite porphyry with potassic alteration (Balsapamba). F – hornblende‐ and biotite‐bearing granodiorite with potassic alteration at Apuela (Cuellaje). G – hornblende granodiorite porphyry at Apuela (Junin) with pervasive potassic altera‐
tion and sericite alteration haloes around quartz‐pyrite veinlets; note that veinlets and their haloes were removed prior to geochemical analysis. H – hornblende granodiorite porphyry at Apuela (Junin) with pervasive potassic, overprinted by phyllic alteration where feldspars are partly replaced by sericite; this type of alteration significantly affects whole‐
rock Sr contents rendering results petrogenetically insignificant; used only for a limited number of Junin porphyry sam‐
ples where other samples were not available. White scale bar is 1 mm for micro‐, and 2 cm for macro‐photographs.
122 123
124 whose roots are exposed in the Eastern Cordillera (Litherland et al. 1994). The forearc sliver and Western Cordillera basement consist of oceanic plateau fragments accreted to the paleo‐
continental margin in the Late Cretaceous and interpreted to form part of the Colombian‐
Caribbean oceanic plateau (CCOP; e.g., Vallejo et al. 2009). Basement units of the Sierra region be‐
tween the Eastern and Western Cordillera are obscured by NVZ and Tertiary arc volcanic prod‐
ucts; they likely consist of a tectonized mélange of oceanic plateau units and Eastern Cordillera basement where the proportion of the latter progressively increases towards the east (Fein‐
inger & Seguin 1983; Litherland et al. 1994; Spik‐
ings et al. 2005; Chapter 4). Southwestern Ecua‐
dor additionally contains a rotated micro‐
continental block known as the El Oro massif whose basement units are of Eastern Cordillera petrogenetic affinity (Litherland et al. 1994). Arc magmatism resulting from the subduction of the Farallon plate below South America was ac‐
tive along the Ecuadorian margin until the Late Jurassic (Litherland et al. 1994). Except for the Mid‐Cretaceous Tangula batholith (Hall & Calle 1982) and few minor backarc intrusions (Barra‐
gan et al. 2005), voluminous Cretaceous arc magmatism has not been identified in Ecuador suggesting a magmatic lull during that period. Following oceanic plateau accretion(s), arc mag‐
matism resumed in the latest Cretaceous, and a continuous Tertiary arc developed along the whole continental margin (Vallejo et al. 2009; Chapter 2). While southern Ecuador is mostly characterized by continental, subaerial arc mag‐
matism throughout the Tertiary, magmatism in northern‐central Ecuador started as a submarine island arc system represented by the Macuchi Unit which was erupted on accreted oceanic pla‐
teau basement; this arc was juxtaposed land‐
wards against a minor subaerial arc system and progressively matured during the Tertiary, culmi‐
nating in the present‐day NVZ arc magmatism on substantially thickened crust (Guillier et al. 2001; Jailliard et al. 2005; Chiaradia 2009; Chiaradia et al. 2009a). Emplacement of multiple intrusions along major structures led to the development of batholith‐size intrusive clusters in northern, cen‐
tral, and southern Ecuador during the Oligocene‐
Miocene, and was accompanied by voluminous arc volcanism (Chapter 2; Fig. 1). Following the Late Cretaceous oceanic plateau accretion, oblique Farallon/Nazca‐South America plate convergence characterized the Ecuadorian margin throughout the Tertiary (Chapter 2). A major geodynamic event affected the Tertiary subduction system with the fragmentation of the Farallon plate, followed by initiation of Cocos‐
Nazca seafloor spreading during the Early Mio‐
cene (Lonsdale 2005; Barckhausen et al. 2008). At the Colombian‐Ecuadorian margin, this led to a change in subducting slab properties from old, cool (Farallon) to young, hot (Nazca) oceanic lithosphere; the Farallon‐Nazca plate boundary is represented by the offshore ENE‐trending Grijal‐
vas scarp progressively propagating southwards along the margin, and currently intersects the Ecuadorian trench at 3°S (Lonsdale 2005). Farallon plate motion reconstructions (Somoza 1998) imply that seamounts formed at the Gala‐
pagos hotspot during the Late Cretaceous to Mid‐
Tertiary did not collide with the Ecuadorian mar‐
gin, situated due east with respect to the hotspot in a present‐day global reference frame; instead, they were subducted at, or docked onto the Pa‐
nama‐Costa Rica margin to the NE (Hoernle et al. 2002). This situation changed fundamentally in the Late Oligocene‐Early Miocene when a major change in Farallon plate motion prior to its fission in the Early Miocene caused Galapagos‐derived seamounts to drift eastwards, resulting in the collision of the Carnegie Ridge seamount chain with the Ecuadorian margin in the Late Miocene where it caused minor shallowing of the subduc‐
tion angle, and eastward arc migration and broa‐
dening (Gutscher et al. 1999; Guillier et al. 2001; Chapters 2, 3). While Gutscher et al. (1999), based on the erroneous assumption of a flat slab geometry below central Ecuador, proposed a slab tear along the projected trace of the Grijalvas scarp below southern‐central Ecuador, the shal‐
low (25‐30°) subduction setting below central Ecuador inferred from high‐resolution seismic studies (Guillier et al. 2001) makes a slab contor‐
tion below southern Ecuador a more likely alter‐
native. A slab contortion is also proposed for the flat‐normal slab transition between northern and central Peru (James & Sacks 1999). The flat slab geometry initiating below northern Peru in the 125
Late Miocene (e.g., James & Sacks 1999) contin‐
ues northwards into southern Ecuador where Late Miocene or younger arc magmatism is con‐
sequently not observed (Chapter 2). Sampling and analytical tech‐
niques Samples for geochemical analyses were collected from mineral exploration drill cores or outcrop exposures of multiple intrusions and volcanic formations listed in Table 1. Sampling did not fol‐
low a systematic grid pattern but reflects drill core distribution as well as outcrop accessibility and suitability. Sampled Quimsacocha volcanics are all related to a single volcanic caldera, whe‐
reas sampling localities for Saraguro Group vol‐
canics show a large geographical spread (c. 100 km) comprising three areas (Chaucha, Cañi‐
capa, Portovelo; Fig. 1) to reflect its widespread distribution in southern‐central Ecuador. Both intrusive centers and Saraguro Group volcanics integrate magmatic events of up to 10‐15 m.y. (the time span for batholith construction in cen‐
tral and northern Ecuador; Chapter 2). To ensure compositionally representative geo‐
chemical analyses, sample quantities typically comprised c. 0.5 or c. 1 kg of material for fine‐ or coarse‐grained samples, respectively. Lower sample quantities were occasionally obtained for drill core samples where sampling material was limited. In areas where rocks were affected by porphyry intrusion‐related hydrothermal sys‐
tems, careful outcrop selection and drill core quick‐logging ensured sampling of least altered material for a given alteration facies. Intense feldspar‐destructive phyllic and argillic alteration was avoided where possible (essentially every‐
where except for some Junin porphyry intru‐
sions). Pervasively veined material was avoided for sampling, and isolated hydrothermal veins and vein alteration haloes were removed by a diamond blade disc saw prior to sample process‐
ing. Samples were cleaned with water, crushed using a steel jaw crusher, and powdered (<70 μm) us‐
ing an agate disc mill, where each step was fol‐
lowed by sample splitting into representative 126 Table 2: Hydrothermal alteration-influenced
compositional variability of exemplary reference lithologies
Elements
Remarks
Low compositional variation
SiO2, TiO2,
Al2O3
scatter < ±10% for all systems
Zr, Hf
scatter < ±10% for all systems
Moderate compositional variation
CaO, P2O5
scatter < ±20% for all systems
Sr
scatter < ±10% for Apuela-Cuellaje, Gaby,
Balsapamba; scatter ~20% (sometimes higher)
for Saraguro
Nb, Ta
scatter < ±20% for Apuela-Cuellaje, Balsapamba, Saraguro; mostly < ±20%, but up to 35%
scatter for Gaby
Y
scatter < ±20% for Gaby; variable (up to 38%)
for Apuela-Cuellaje, Balsapamba, Saraguro
Sc
scatter < ±20% for Apuela-Cuellaje, Balsapamba, Saraguro; slightly higher scatter (up to
26%) for Gaby
Th
scatter < ±20% for all systems
V
scatter < ±10% for Apuela-Cuellaje, Balsapamba, Saraguro; systematic bias at Gaby
caused by reference sample composition, otherwise scatter at Gaby would be < ±20%
High compositional variation
Fe2O3, MgO scatter < ±20% for Apuela-Cuellaje, Balsapamba, Saraguro; highly variable scatter for
Gaby (up to 52%)
Na2O
scatter < ±10% for Apuela-Cuellaje, Balsapamba, Saraguro; highly variable scatter for
Gaby
K2O
highly variable scatter for all systems
Cs, Rb, Ba
highly variable scatter for all systems; at Gaby
correlated with K2O (potassic alt.)
U
highly variable scatter for all systems
Cr, Ni
highly variable scatter for all systems
REE
highly variable scatter for all systems; LREE
decoupled and La/Yb ratios potentially inaccurate; coupled behavior within MREE & HREE
groups such that Eu/Eu* and Dy/Yb are potentially accurate
Main alteration types for reference centers where isocons
were constructed
Apuela
(Cuellaje)
potassic, propylitic
Balsapamba potassic, propylitic
Gaby
sodic-calcic, potassic, propylitic
Saraguro
propylitic
Single-analysis outliers are not noted, but might be significant;
refer to Table A2 (Appendix). Note that feldspar phenocrystdestructive alteration was generally avoided as far as possible,
but is unavoidable for some samples at Junin and Chaucha; Sr
scatter is expected to increase for these centers.
proportions to reduce the sample quantity for subsequent analytical steps. Reconnaissance XRF analysis for major and trace elements (10‐15 g/sample) for a total of 139 samples was carried out during the 2006‐2008 period at the Institute of Mineralogy and Geochemistry, University of Lausanne. XRF measurements were done on pressed powder pellets or fused glass beads Figure 3: Isocon plots of representative Tertiary granitoids and volcanics in Ecuador, constructed after Grant (1986) assuming constant concentrations of Al2O3 (the least alteration‐affected major element oxide) between altered and least‐altered reference sample. Elemental scatter outside gray area (± 10% analytical uncertainty) reflects alteration‐
induced compositional variation (± lithologic compositional heterogeneities) of altered samples with respect to a least‐
altered reference sample of the same lithology. Note that displayed isocons correspond to average compositions of all
compared lithologies and serve only for illustrative purposes; to quantify alteration effects individual isocons were cal‐
culated for each sample and considered for relative concentrations shown in the Appendix (Tab. A2). Isocon slopes vary
according to dehydration‐induced mass loss effects. Alteration acronym key: K – potassic; Na‐Ca – sodic‐calcic; propyl. – propylitic. Feldspar‐destructive phyllic or argillic alteration is expected to increase Sr mobility but was avoided for
most samples. See Seedorff et al. (2005) for definitions of alteration assemblages. Refer to Appendix for discussion of alteration effects, and to Table 2 for a summary of expected alteration effects on sample compositions. 127
fluxed with Li2B4O7 using a Philips PW 2400 ana‐
lyzer. Data accuracy, precision, and reproducibil‐
ity were controlled using a number of natural and synthetic international or in‐house standard ma‐
terials (BHVO, EMU3.14, QLO, QTW, MFTH‐1, NIM‐G, SDC‐1, SY‐2) which were selected accord‐
ing to sample composition and analytical meas‐
urement program. Expected 2σ uncertainties from repeated standard measurements are 2‐7% for major elements and up to 10% for trace ele‐
ments. Combined microscopic analysis and XRF result screening served to select only least altered sam‐
ples in a given intrusive suite or volcanic forma‐
tion for subsequent laser‐ablation inductively‐
coupled plasma mass spectrometry (LA‐ICP‐MS) analysis. LA‐ICP‐MS analysis of glass bead frag‐
ments recovered from XRF analysis (n = 97) was carried out at the Institute of Mineralogy and Geochemistry, University of Lausanne, using a 193 nm Lambda Physik Excimer Laser system as‐
sociated with a Perkin‐Elmer ELAN 6100 DRC quadrupole ICP mass spectrometer. The Laser system was operated at 10 Hz frequency using 140 or 160 mJ output energy and a 120 μm beam diameter. Background measurement for c. 90 s was followed by 30‐40 s sample ablation with 3‐4 ablation pits for each sample. Used internal stan‐
dard elements were CaO (for sample CaO >1 wt.%) or Al2O3 (for sample CaO <1 wt.%), refer‐
enced to NIST SRM‐610 and SRM‐612 standard materials and sample major element composi‐
tions measured by XRF. Off‐line data reduction including automatic spike correction used the Matlab‐based SILLS codec (Guillong et al. 2008). Trace elements measured by XRF and selected for data presentation in this study comprise Sr, Y, Pb, V, Cr, Cu, Zn, and Ga, whereas for all other trace elements LA‐ICP‐MS results were preferred. X‐ray fluorescence and LA‐ICP‐MS trace element analy‐
ses generally agree within error, except for Nb, where a systematic bias of elevated contents for XRF measurements was detected, and Zr, where XRF results scattered outside of the expected ±10% analytical error limit. Results Combined whole‐rock XRF and LA‐ICP‐MS major and trace element compositions of the complete 128 dataset are listed in the Appendix (Tab. A1). Samples are grouped according to localities (in‐
trusions) and stratigraphic units (volcanics; Tab. 1). Considering the constraints on sampling pro‐
cedures described above, it is important to em‐
phasize that samples within single and between different sample groups are generally not consid‐
ered as cogenetic. Apart from sharing an overall similar petrogenesis in terms of Late Oligocene to Late Miocene arc magmatism, their chemical compositions are therefore not systematically related to each other by a single specific petro‐
genetic process (such as magma mixing or AFC, i.e., assimilation and fractional crystallization). Rock alteration and element mobility Depending on their spatial and temporal prox‐
imity with respect to centers of porphyry‐related hydrothermal systems, most rocks display vari‐
able degrees of high‐ to low‐temperature altera‐
tion comprising potassic, sodic‐calcic or calcic‐
sodic, phyllic, (advanced) argillic and propylitic alteration assemblages (e.g., Seedorff et al. 2005; Fig. 2). Effects of hydrothermal alteration on rock chemistry in our dataset are summarized in Table 2 and Figure 3, and are further discussed in the Appendix. Element mobility and redistribution due to hy‐
drothermal alteration commonly strongly affects large ion lithophile elements (LILE), whereas high field strength elements (HFSE) are less affected. As shown above (Fig. 3, Tab. 2; see also discus‐
sion in the Appendix) certain LILE (e.g., Cs, Rb, as well as K as major element) are strongly affected by hydrothermal alteration and are thus not suited as petrologic tracers, whereas most HFSE and Sr (although a LILE, and often considered as mobile during hydrothermal alteration) tend to stay relatively immobile in our dataset. Light, mid‐, and heavy rare earth elements (LREE, MREE and HREE, respectively) do not consistently show chemically coupled behavior during hydrothermal alteration, whereas alteration‐induced intra‐
MREE and HREE scatter appears to be coupled in our reference samples (Fig. 3, Tab. 2; Appendix). Consequently, while chondrite‐normalized REE distribution patterns can be used for a qualitative petrogenetic discussion, LREE/HREE ratios such as La/Yb are potentially inaccurate; intra‐HREE and MREE/HREE ratios such as Dy/Yb and Sm/Dy, on the other hand, are more likely to reflect petro‐
genetically significant, rather than alteration‐
biased values. We emphasize that these observa‐
tions reflect our sample selection according to alteration mineralogy, and cannot be generalized for other datasets. Throughout this article we avoid using any chemical element for petroge‐
netic discussions whose concentration is consid‐
ered as significantly affected by alteration in one of our reference magmatic centers (Tab. 2, Ap‐
pendix; some of these elements may be immobile in several of the investigated magmatic centers though). Rock petrography Geologic features of sampled magmatic centers are summarized in Table 1. The igneous (as op‐
posed to alteration‐induced hydrothermal) min‐
eralogical inventory of most phaneritic intrusive rocks comprises plagioclase, hornblende, quartz, and biotite with accessory opaque minerals Figure 4: Rock petrographic classification of Late Tertiary granitoids and volcanics in Ecuador. Upper left: plutonic rock classification by normative quartz‐alkali feldspar‐plagioclase (QAP) proportions (Gillespie & Styles 1999) using projec‐
tion parameters of Le Maitre (1976); due to alteration‐induced Ca depletion, normative albite was used instead of an‐
orthite to calculate the plagioclase component of Junin porphyry intrusions. CIPW norms were calculated assuming Fe2+
= 0.7 ∑ Fe. Upper right: alumina saturation index classification of plutonic rocks (Maniar & Piccoli 1989). Lower left:
Total alkali‐silica classification for plutonic and volcanic rocks (Le Bas et al. 1986). Lower right: SiO2 vs. Zr/TiO2 / 10,000 classification for plutonic and volcanic rocks (Winchester & Floyd 1977); note that using Nb/Zr instead of SiO2 as differ‐
entiation index produces systematic inaccuracies for rock classification in our dataset and is therefore not applied. 129
(mostly magnetite), apatite, and zircon; most highly differentiated intrusions additionally con‐
tain minor alkali‐feldspar and accessory titanite. Most of these intrusions mineralogically classify as tonalite, granodiorite, or quartz‐diorite (Fig. 2). In decreasing order of abundance, phenocryst modes of porphyry intrusions mainly comprise plagioclase, ±hornblende, ±quartz, ±biotite, with quartz commonly displaying rounded, resorbed grain margins. Sampled Quimsacocha andesite lava flows and dacite domes contain plagioclase and hornblende phenocrysts (the latter often opacitized), frequently embedded in a glassy, non‐devitrified matrix; clinopyroxene occurs as additional phenocryst in andesites. Saraguro vol‐
canics comprise andesitic, dacitic, and rhyolitic flows and subordinate tuffs where main phenoc‐
ryst assemblages are plagioclase, ±hornblende, ±clinopyroxene, plus quartz with accessory zircon for dacitic‐rhyolitic compositions. Chemical classification plots based on normative mineral proportions of quartz, alkali‐feldspar, and plagioclase (QAP), total alkali versus silica con‐
tents, or ratios of immobile trace elements (Fig. 4) yield consistent results, with most samples classifying as andesitic to rhyo‐dacitic in composi‐
tion, or as tonalite, granodiorite, or monzogranite in normative QAP proportions. Based on com‐
parison with visually estimated mineral modes, the alkali‐feldspar component calculated from normative mineral proportions in Figure 4 seems to be slightly, and in the case of Zaruma‐
Portovelo samples moderately, overestimated due to hydrothermal alteration. Most intrusions straddle the metaluminous‐peraluminous border (with a tendency to plot in the peraluminous field) with only Gaby classifying as entirely meta‐
luminous (Fig. 4). Major element contents Most samples define broadly continuous distribu‐
tion patterns in Harker diagrams (Fig. 5) where TiO2, Fe2O3 and CaO steadily decrease with SiO2, i.e., they show gross compatible behavior, whereas K2O and Na2O mainly display incompati‐
ble behavior or constant values with increasing SiO2. Al2O3 and CaO/Al2O3 distribution trends di‐
verge and show either a broad decrease or con‐
130 stant values with increasing SiO2. Second‐order scatter of major elements partly reflects hydro‐
thermal alteration (especially for K2O). Most samples seem to classify as low‐K series or strad‐
dle the low‐ to medium‐K series border (Fig. 5), but rock alteration precludes a detailed classifica‐
tion. Except for Quimsacocha dacites, all mag‐
matic centers plot in the calc‐alkaline field of Mi‐
yashiro (1974). The broad similarity of major element trends displayed by the whole dataset suggests that similar processes of magma evolu‐
tion operate in all investigated Oligocene‐
Miocene magmatic centers. Trace element contents Where appropriate, trace element Harker dia‐
grams (Figs. 6, 7) are plotted together with the compositional fields of the present‐day NVZ main and frontal arc (Chiaradia et al. 2009a), and Eo‐
cene lavas of the Macuchi Unit (Chiaradia 2009). These references were chosen for increased in‐
ternal analytical consistency with our dataset as they were measured in the same laboratory; more compositional NVZ data are available (e.g., Bourdon et al. 2003; Samaniego et al. 2005; Bry‐
ant et al. 2006), and mostly overlap with the given reference fields. In primitive mantle‐normalized spidergrams all Late Tertiary samples display negative Nb‐Ta anomalies relative to LILE indicative of a slab flu‐
id‐metasomatized mantle wedge as the magma source (Fig. 8). With the exception of few Sara‐
guro Group samples, trace elements incompati‐
ble in basaltic melts (Th, U, Zr, Hf) consistently plot below the present‐day main arc, but overlap with NVZ frontal arc compositions (Fig. 6). Within a given Late Tertiary magmatic center, these trace elements for the most part either stay con‐
stant or show decreasing contents at SiO2 >65 wt.% possibly reflecting the influence of ac‐
cessory phase fractionation which only become stabilized in relatively silicic melts (e.g., Hoskins et al. 2000). Trace elements compatible in basal‐
tic melts show petrologically correlated behavior with the previous group in that they consistently plot above, or overlap with, frontal (and main) arc NVZ magmas; overall, basalt‐compatible trace Figure 5: Major element oxide vs. SiO2 concentrations of investigated samples. Low‐K, medium‐K, and high‐K classifica‐
tion from Gill (1981). Tholeiitic vs. calc‐alkaline dividing line from Miyashiro (1974). 131
element contents (e.g., Y, Yb, Sc) decrease with increasing SiO2. Strontium generally shows broadly constant con‐
centrations around 300 ppm for the 55‐65 wt.% SiO2 interval and major scatter (100 to c. 800 ppm) for more evolved compositions, al‐
though average Sr contents do not shift system‐
atically for compositions >65 wt.% SiO2 if Quim‐
sacocha is excluded (Fig. 6). With the exception of Quimsacocha, Sr contents of Late Tertiary arc magmas between 55‐65 wt.% SiO2 overlap with, or slightly exceed the upper compositional range defined by the less differentiated Macuchi ba‐
salts‐andesites, but consistently plot below both main and frontal arc NVZ magmas, although no data for <59 wt.% SiO2 exist for the latter (Fig. 6). Trace element ratios (Fig. 7) illustrate the signifi‐
cance of these general compositional features for the development of adakite‐like chemical signa‐
tures: the majority of Late Tertiary Western Cor‐
dillera granitoids (Apuela‐Junin, Balsapamba‐
Telimbela, Chaucha) qualify as adakite‐like in a Sr/Y vs. Y discrimination plot (we do not use the additional La/Yb vs. Yb adakite discrimination plot because of potential LREE/HREE decoupling dur‐
ing alteration as mentioned above) due to strongly depleted Y contents (<10 ppm) and de‐
spite only moderately high (mostly sub‐adakite‐
like, i.e., <400 ppm) Sr contents. A Sr/Y vs. SiO2 plot demonstrates that most present‐day NVZ magmas acquire adakite‐like Sr/Y ratios during relatively early differentiation stages, whereas most Oligocene‐Miocene magmas (excluding Quimsacocha) only display clear adakite‐like fea‐
tures for highly differentiated (>65 wt.% SiO2) compositions (Fig. 7; note, however, that for the Apuela batholith and Junin porphyries Sr and Sr/Y partly decrease with increasing SiO2 in the 65‐70 wt.% SiO2 interval; this might be an alteration effect). Quimsacocha, in contrast to other Oligo‐
cene‐Miocene magmatic centers, tends to over‐
lap with NVZ magmas, although it is depleted in Y with respect to the latter (Figs. 6, 7). These sys‐
tematic differences imply that Oligocene‐
Miocene magmas were either derived from a source with a different trace element composi‐
tion than present‐day main arc NVZ magmas, or/and that the latter acquired their surplus in Sr (and other basalt‐incompatible elements) during initial stages of basalt differentiation in a MASH 132 or hot zone (Hildreth & Moorbath 1988; Annen et al. 2006) through differing AFC ± mixing proc‐
esses compared to Oligocene‐Miocene magmas. These contrasting magma evolution trends can be bracketed in time between the Quimsacocha volcano (7.1 Ma; Chapter 2), whose composi‐
tional trend largely mirrors present‐day NVZ magmas, and the Chaucha batholith (14.8 Ma; Chapter 2). Younger porphyry intrusions at Junin and Chaucha (9.0 and 9.8 Ma; Chapter 2) are too highly differentiated to confidently predict their compositional behavior in the differentiation in‐
terval relevant for NVZ magmas (<65 wt.% SiO2). Chondrite‐normalized REE plots (Fig. 8) display moderate to strong LREE enrichments over HREE, except for Gaby, where relatively flat REE pat‐
terns suggest parental magma differentiation was not driven by strongly HREE/LREE‐fractionating mineral phases. Most REE plots lack negative Eu anomalies indicative of plagioclase fractionation (Fig. 8; see also relatively constant Eu/Eu* range at c. 1.0±0.4 in Fig. 7); however, negative Eu anomalies do characterize a number of mainly southern Ecuadorian magmatic centers compris‐
ing Saraguro Group volcanics at Chaucha and Por‐
tovelo‐Zaruma (but not at Cañicapa and Tres Chorreras), several intrusions of the Cangrejos‐
Zaruma intrusive belt (Cangrejos porphyries; in‐
trusions north of Zaruma), and, partly, the Telim‐
bela batholith. Positive Eu anomalies (see discus‐
sion below for their petrogenetic interpretation) are displayed by a fraction of Balsapamba batho‐
lith samples and the Gur‐Gur porphyry intrusion at Chaucha. Several magmatic centers are strongly depleted in HREE and display concave‐
upwards HREE patterns typical of amphibole frac‐
tionation. Chondrite‐normalized Dy/Yb ratios may be used as petrologic fingerprints for garnet or amphibole fractionation (leading to strongly increasing or slightly decreasing trends with increasing SiO2; Davidson et al. 2007). Individual Quaternary NVZ volcanic centers show Dy/Yb and Nb/Ta trends negatively correlated with SiO2 indicative of am‐
phibole fractionation (Chiaradia et al. 2009a). No such systematic trends can be clearly discerned for Oligocene‐Miocene arc magmas; Nb/Ta ratios are subchondritic and, mostly, sub‐primitive mantle (Fig. 7). Figure 6: Trace element vs. SiO2 concentrations of investigated samples; two columns arranged in order of downwards decreasing trace element incompatibility in basaltic melts. Macuchi reference fields from Chiaradia (2009); NVZ main and frontal arc from Chiaradia et al. (2009a). 133
The lack of systematic Dy/Yb and Nb/Ta distribu‐
tion trends is not surprising given that most sam‐
ples do not define cogenetic suites, such that dif‐
ferentiation trends of individual magmatic sys‐
tems are not visible, and any long‐term bulk dis‐
tribution trends might additionally become blurred by the somewhat contrasting, superpos‐
ing effects of amphibole and garnet (or other in‐
tra‐HREE‐fractionating phases; e.g., Davidson et al. 2007) fractionation. In contrast, Sm/Dy ratios will increase in response to both amphibole and garnet fractionation and are thus supposed to show a more homogeneous distribution pattern. This is evidenced in Figure 7 where chondrite‐
normalized Sm/Dy ratios show an overall steady increase with SiO2, albeit showing pronounced variations between different magmatic centers for a given differentiation stage. With few excep‐
tions, Sr/Y and Sm/Dy show a broad positive cor‐
relation (Fig. 7). Rare earth element distribution patterns Four major groups of variable REE distribution patterns can be distinguished for the Late Terti‐
ary Ecuadorian intrusions and volcanic rocks in‐
vestigated in this study (Fig. 8; Tab. 3). These groups mainly differ in their ways of HREE frac‐
tionation and in the occurrence or absence of negative Eu anomalies. The facts that (1) no Ce anomalies are observed, and (2) the main vari‐
ability in the REE pattern shape lies in the MREE and HREE instead of the LREE (Fig. 8) have been used elsewhere to argue for the petrogenetic significance of REE patterns of altered rocks (Kay et al. 2005). Group 1: REE patterns characterized by negative Eu anomalies and without strong HREE depletion This group mainly comprises volcanics and phan‐
eritic‐porphyritic intrusions in southern Ecuador (Cangrejos‐Zaruma intrusive belt, El Mozo, Sara‐
guro Group at Chaucha and Portovelo). Addition‐
ally, few samples of the Apuela and the Balsa‐
pamba and Telimbela batholiths share the same characteristics. Rocks of this group uniformly dis‐
play minor‐moderate negative Eu anomalies in 134 their REE patterns, and have SiO2 contents in the range of 53‐66 wt.% (except for a single higher differentiated Apuela sample). For the most part, these plutonic and volcanic rocks show relatively flat MREE‐HREE patterns with HREE concentra‐
tions ≥10x chondrite; slightly fractionated or U‐
shaped HREE patterns are rare. All intrusions are characterized by high modal proportions of pla‐
gioclase, and their REE distribution patterns sug‐
gest that their parental melts fractionated plagio‐
clase. Although amphibole is present in signifi‐
cant modal proportions in several intrusions, sig‐
nificant amphibole fractionation did not take place. In H2O‐saturated experimental runs of ba‐
saltic‐andesitic bulk compositions buffered at NNO, plagioclase appears earlier on the liquidus than amphibole only at pressures <0.4 GPa (Grove et al. 2003; the maximum pressure might increase towards higher degrees of melt H2O‐
undersaturation) suggesting that parental melts of intrusions and volcanics of this group under‐
went a significant evolution step at shallow crus‐
tal levels. Prior to their shallow crustal evolution, deeper parental melt evolution in a MASH/hot zone mostly did not involve significant amphibole (or garnet) fractionation. The latter condition might apply for primitive melt evolution in a hot zone at the base of a relatively thin crust (or a mid‐crustal hot zone in a thicker crust) where H2O‐undersaturated parental melts would mainly crystallize pyroxene instead of amphibole (Müntener et al. 2001). Group 2: REE patterns without nega‐
tive Eu anomalies or strong HREE de‐
pletion This group comprises variably differentiated (52‐
71 wt.% SiO2) intrusions (Gaby, Chaucha, Cangre‐
jos, Zaruma) and Saraguro Group volcanics (at Cañicapa and Tres Chorreras) which neither show significant negative Eu anomalies nor strongly depleted HREE contents, although slightly con‐
cave‐upwards MREE‐HREE patterns can be ob‐
served for some samples. Parental melts of this group probably share a similar deep‐ to mid‐
crustal evolutionary history with the previous group in that parental melt evolution at depth was dominated by pyroxene fractionation, al‐
though it might have included minor amphibole Figure 7: Trace element ratios vs. SiO2 contents of investigated samples. Reference fields as in Fig. 6 where applicable.
135
fractionation in some cases where more hydrous melts were involved (e.g., the Portovelo porphyry intrusions; Fig. 8). Intra‐HREE fractionation due to garnet fractionation cannot be observed in any of these magmas. In contrast to the previous group, the absence of negative Eu anomalies is indica‐
tive of the absence of significant plagioclase frac‐
tionation and thus suggests limited shallow crustal open‐system magma evolution. Alterna‐
tively, plagioclase fractionation took place in highly oxidized melts where Eu incompatibility for plagioclase increases (e.g., Rollinson 1993). Group 3: REE patterns with strong HREE depletion and concave‐upwards to flat HREE distribution Significant HREE depletion (<10x chondritic val‐
ues) combined with mostly concave‐upwards or minor relatively flat HREE patterns is observed for mainly porphyritic, plus some phaneritic intru‐
sions at Chaucha (Gur‐Gur), Balsapamba‐
Telimbela, and Apuela‐Junin (Fig. 8). All intrusive rocks of this group are highly differentiated (64‐
73 wt.% SiO2). In several cases strong HREE frac‐
tionation coincides with the development of a minor‐moderate positive Eu anomaly. The latter samples are characterized by moderate Sr con‐
tents (300‐400 ppm; see discussion below), and petrographic studies indicate that they do not represent plagioclase cumulates, although modal proportions of plagioclase are relatively high. While individual mineral phases (e.g., amphibole, titanite, plagioclase, garnet, zircon) impose char‐
acteristic fractionation signatures on REE pat‐
terns, their integrated (not necessarily concomi‐
tant) fractionation effects might in part produce mutual offsets of otherwise typical REE character‐
istics. This might be of relevance for our dataset, and the following considerations apply in this context: 
136 The strong HREE depletion displayed by samples of this group (and, possibly, their partly positive Eu anomalies) are indicative of amphibole and/or titanite fractionation (or equilibration with a restite bearing these minerals; e.g., Davidson et al. 2007). Given the highly differentiated sample compositions, liquid‐amphibole HREE parti‐
tion coefficients will be fairly high (Bach‐
mann et al. 2005) such that small amounts of amphibole fractionation, or trace amounts of titanite fractionation (Bach‐
mann et al. 2005; Glazner et al. 2008) both succeed in explaining strong HREE deple‐
tion. HREE fractionation patterns do not suggest significant garnet fractiona‐
tion/restite equilibration (but see below). 
Negative DEu anomalies for titanite or am‐
phibole and the resulting positive Eu ano‐
malies in REE patterns may partly offset va‐
riably (depending on the prevailing fO2) negative Eu anomalies imposed on REE pat‐
terns by plagioclase fractionation; depend‐
ing on P‐T conditions, melt composition, and fO2 an integrated fractionating mineral assemblage consisting of plagioclase + am‐
phibole + trace amounts (<1%) of titanite can produce relatively smooth MREE pat‐
terns without negative Eu anomalies (Glazner et al. 2008). Smooth LREE‐MREE patterns, and/or positive Eu anomalies (which are probably not attributable to plagioclase accumulation in our samples) observed for some intrusions of this group illustrate this effect (Fig. 8). Minor plagio‐
clase fractionation, although it is not obvi‐
ously manifested as negative Eu anomalies in the REE patterns of this group, thus can‐
not be ruled out for parental melts of intru‐
sions of this group. However, as shown be‐
low, non‐depleted Sr contents of these samples indicate that plagioclase fractiona‐
tion was not significant. The absence of concave‐upwards MREE‐HREE patterns characteristic of amphibole (and ti‐
tanite) fractionation for some intrusions might signal minor additional fractionation/restite equilibration with a HREE‐fractionating residual mineral with DDy < DYb such as garnet (or zircon), prior to, or concomitant with amphibole (or ti‐
tanite) fractionation. Alonso‐Perez et al. (2009) show that garnet and amphibole are both the first liquidus phases in H2O‐rich andesitic melts at high pressures; their modal proportions vary from garnet‐dominated (at 1.2 GPa) to amphi‐
bole‐dominated (at 0.8 GPa). While both minerals drive bulk HREE depletion, their integrated frac‐
tionation effects (or restite equilibration) might Figure 8: C1 chondrite‐normalized REE diagrams and primitive mantle‐normalized spidergrams of investigated samples.
Note that LILE scatter in spidergrams is an effect of hydrothermal alteration, such that these elements are of limited petrogenetic relevance. Normalization values from Sun & McDonough (1989) 137
Figure 8 (continued) partly mask their characteristic fingerprints on intra‐HREE fractionation (e.g., Dy/Yb). Most parental melts of this group of intrusions are thus inferred to have undergone significant amphibole (± minor titanite; see below) frac‐
tionation. Additional effects from integrated, relatively minor plagioclase or garnet fractiona‐
tion, as well as from pyroxene fractionation (as the groups above) may be variably superposed on the observed patterns, but do not strongly influence REE distribution characteristics. 138 Group 4: REE patterns with strong HREE depletion and a negative HREE slope This group, comprising the Quimsacocha volcanic center and some Junin porphyry intrusions, is characterized by significant HREE depletion and steadily decreasing MREE/HREE ratios (Fig. 8). Negative Eu anomalies are not observed. A single sample of a Balsapamba porphyry intrusion shows similarly strong HREE fractionation, but displays a minor negative Eu anomaly; it is there‐
fore included in group 1 above, but the following considerations might apply for the deep crustal petrogenesis of this sample. Quimsacocha dacites show flatter bulk REE patterns than Quimsacocha andesites implying that they cannot directly rep‐
resent derivative liquids from andesite differen‐
tiation; dacite REE patterns resemble Quimsaco‐
cha andesites, but both LREE/HREE and intra‐
HREE fractionation is less pronounced. Junin por‐
phyry intrusions show similar intra‐HREE frac‐
tionation as Quimsacocha dacites, but their LREE/HREE fractionation is slightly less pro‐
nounced; the latter might be an alteration effect as these samples often show pervasive phyllic alteration, in agreement with the general notion that LREE are more mobile in fluids than HREE (see above and Appendix). The strongly fractionated intra‐HREE patterns combined with moderate‐strong LREE/HREE frac‐
tionation observed in this group closely resemble some Quaternary arc volcanoes such as Young Chacana, whose parental melts are inferred to evolve through combined amphibole, clinopyrox‐
ene, and garnet fractionation/restite equilibra‐
tion (Chiaradia et al. 2009a). These minerals may be stabilized in variably H2O‐rich melts processed in deep to mid‐crustal hot zones (e.g., at 0.8‐1.2 GPa; Alonso‐Perez et al. 2009). The similarities in REE characteristics between some present‐day NVZ magmas and the Late Miocene Quimsacocha and some Junin porphyry magmas suggest that they might share a similar petrogenesis. In the latter case, slab melting (Beate et al. 2001) would not be required to explain intra‐HREE fractiona‐
tion (and other adakite‐like geochemical fea‐
tures) of the Quimsacocha volcanic center. Adakite‐like features of Late Ter‐
tiary Ecuadorian arc magmas When discussing adakite‐like features, we focus in the following on Sr and Y concentrations and the Sr/Y ratio, the latter often representing the most distinctive discrimination criteria for ada‐
kite‐like magmatism at convergent margins (e.g., Tulloch & Kimborough 2003). Furthermore, for the majority of samples in our dataset (see dis‐
cussion above and Appendix) the Sr/Y ratio is more robust with respect to hydrothermal altera‐
tion than other adakite discrimination criteria (e.g., La/Yb). We correlate our observations on Sr and Y distributions with petrogenetic constraints obtained from REE patterns discussed above. Strontium contents Despite a lack of negative Eu anomalies in the majority of Late Oligocene to Late Miocene intru‐
sions and volcanics suggesting limited plagioclase fractionation, whole‐rock Sr contents are mostly relatively low: most samples show Sr contents around the 348 ppm value of the average lower crust of Rudnick & Gao (2003) and <410 ppm (av‐
erage andesite of Gill 1981); except for Quimsa‐
cocha, most samples invariably plot below pre‐
sent‐day main arc values at a given SiO2 content, and there is a broad bulk increase in Sr concen‐
trations through time (Fig. 9). These low Sr con‐
tents might reflect source and/or a crustal magma differentiation effects. Considering the former, the highly variable geodynamic regime at the northern Andean margin in the Late Tertiary makes source controls on the Sr budget of arc magmas by processes operating in the mantle wedge likely (e.g., variable degrees of mantle wedge contamination/metasomatization, or par‐
tial melting). These effects cannot be evaluated in this study, however, as our dataset mainly com‐
prises highly differentiated samples. In the fol‐
lowing, we therefore focus on discussing crustal magma differentiation. Mid‐ to shallow crustal magma differentiation (inside the stability field of plagioclase) Low Sr contents of Late Oligocene to Miocene samples can in part be attributed to shallow crus‐
tal magma differentiation following an initial dif‐
ferentiation step from mantle‐derived basaltic to andesitic‐dacitic compositions in deep to mid‐
crustal MASH/hot zones. This is the case for a group of Miocene intrusions and volcanics (group 1 in the previous section) characterized by nega‐
tive Eu anomalies indicative of plagioclase frac‐
tionation. Significant Sr depletion by extended mid‐ to shallow crustal plagioclase fractionation might be considered unlikely for a number of Late Oligocene to Mid‐Miocene magmatic cen‐
ters (group 2 in the previous section) unless melts were highly oxidized such that negative Eu anomalies were not produced. The latter is pos‐
sibly supported by broadly similar Sr contents of groups 1 and 2 at similar degrees of differentia‐
139
tion (Tab. 3) despite significant variations in the magnitude of Eu anomalies. Overall, Sr contents of group 3 are slightly higher than Sr contents of groups 1 and 2 (Tab. 3). Pla‐
gioclase fractionation hence was probably minor in parental melts of group 3 intrusions, as also supported by their REE patterns (Fig. 8). In addi‐
tion, if parental melts of these intrusions were H2O‐rich, as suggested by the amphibole imprint on their REE patterns, plagioclase might be anor‐
thite‐rich. As shown by Blundy & Wood (1991), the plagioclase‐melt partitioning coefficient for Sr decreases with increasing molar fractions of an‐
orthite in plagioclase. Consequently, residual melt Sr depletion by (minor) plagioclase frac‐
tionation would be less pronounced. Of further significance might be the tendency of plutonic rocks to be slightly depleted in incom‐
patible elements with respect to cogenetic vol‐
canic rocks (Bachmann et al. 2007). However, this effect should be more pronounced for more in‐
compatible elements such as Th or U (compared to Sr incompatibility), opposite to what is ob‐
served (Fig. 6). Furthermore, volcanic rocks of the Saraguro Group have similar low Sr contents as coeval intrusive rocks (Fig. 6) such that a system‐
atic compositional difference between volcanic and plutonic rocks is an unlikely cause for the observed variations in Sr contents. Deep to mid‐crustal magma differentiation (out‐
side the stability field of plagioclase) Melts extracted from a MASH/hot zone at the base of a thick (c. 50 km) crust variably equili‐
brate with mineral phases of basal crustal litholo‐
gies, e.g., pyroxene or garnet granulites; crustal assimilation or mixing with crustal partial melts of these lithologies results in elevated melt Sr con‐
tents and plagioclase‐diminished crustal residues where the latter effect increases with crustal thickness (Hildreth & Moorbath 1988). The over‐
all crustal thickness in Ecuador during the Oligo‐
cene‐Miocene was lower than at the present day (Jaillard et al. 2005) such that the bulk increase in Sr concentrations from Late Tertiary to Quater‐
nary arc magmas (Fig. 9) might reflect stepwise or progressive crustal thickening. Additionally, in‐
creasing high‐pressure magma differentiation could also be envisaged by downwards migration 140 of the focus of crustal hot zone magmatism (e.g., Mamani et al. 2010). Crustal basement composition Basement (assimilant) composition represents a further potential control factor on arc magma Sr budgets. Mature arc systems (i.e., the Late Mio‐
cene‐Quaternary Ecuadorian arc) include arc in‐
trusive roots which are already enriched in in‐
compatible elements and thus have a higher po‐
tential to contaminate arc magmas of a given starting composition in incompatible elements than a less mature arc system (i.e., the Oligocene to Early Miocene Ecuadorian arc; Hildreth & Moorbath 1988; see also Davidson et al. 1987). However, while this could drive bulk incompati‐
ble element enrichment in Late Miocene‐
Quaternary arc magmas, it would not explain se‐
lective Sr enrichment. Strontium contents of CCOP units as potential assimilant are generally low (typically 11‐267 ppm Sr; e.g., Mamberti et al. 2003) compared to average lower continental crust values (348 ppm Sr; Rudnick & Gao 2003); in contrast, CCOP whole rock Sr contents reported by Allibon et al. (2008) are heterogeneous and extend to significantly higher values (98‐1123 ppm Sr; possibly in part reflecting alteration‐induced increases), such that partial melting and mixing of high‐Sr oceanic pla‐
teau lithologies with mantle‐derived arc magmas in a lower crustal MASH/hot zone might contrib‐
ute to producing derivative liquids with Sr con‐
tents typical for NVZ volcanics (400 to >800 ppm Sr; e.g., Chiaradia et al. 2009a). Significant Late Miocene eastward arc broaden‐
ing (Chapter 2) likely added further lithologies as potential assimilants at depth (whose geochemi‐
cal composition is unconstrained, but might be higher in Sr than oceanic plateau units), and fur‐
ther caused arc magmatism to migrate towards the region of maximum crustal thickness in across‐arc dimension (Jaillard et al. 2005). There‐
fore, the Late Miocene arc migration might con‐
tribute to increased Sr contents of NVZ and Quimsacocha volcanics, relative to Oligocene to Mid‐Miocene Ecuadorian arc magmas. Yttrium contents Maximum and minimum Y contents of Late Oli‐
gocene‐Miocene arc magmas show a broad de‐
crease from 24 to 9 Ma (if the El Mozo intrusions are excluded; Fig. 9). Extreme Y depletion (<10 ppm) can be observed for parts of the Early Miocene Balsapamba‐Telimbela batholith, and the Mid‐ to Late Miocene Chaucha and Apuela‐
Junin batholiths. It includes both phaneritic and porphyritic intrusive rock facies, and correlates with strong HREE depletion (groups 3 and 4 in previous section). Strong Y depletion furthermore affects the Late Miocene Quimsacocha volcanic center. Yttrium depletion in arc magmas has been corre‐
lated with increasing crustal thickness as the lat‐
ter favors the stability of minerals with a strong affinity for Y (in particular, garnet; Hildreth & Moorbath 1988; Richards & Kerrich 2007). The broad correlations of Sr, Y, and Sm/Dy distribu‐
tion trends through time (Fig. 9) are in agreement with progressively increasing high‐pressure crustal magma differentiation by crustal thicken‐
ing. The latter does not apply for Y‐depleted in‐
trusions at Balsapamba‐Telimbela, however, as the crustal thickness in the Early Miocene was significantly below the present‐day crustal thick‐
ness in Ecuador (Jaillard et al. 2005). A possible explanation for these compositional anomalies is a more prominent role of accessory titanite fractionation which might also occur at low pressures (see below). Alternatively, the ir‐
regular geochemical signature of Balsapamba‐
Telimbela might be related to anomalies in the subducting lithosphere such as oceanic fracture zones. The latter may lead to a locally increased volatile flux into the mantle wedge giving rise to unusually H2O‐rich arc magmas where amphibole and garnet stability is increased (e.g., Rodriguez et al. 2007). While Kay et al. (2005) note that lo‐
cal peaks in mantle wedge contamination by subduction erosion may also explain the irregular occurrence of adakite‐like arc magmas, this is rather unlikely for Balsapamba‐Telimbela, as LILE anomalies (parallel to Y and HREE depletion) are not observed, and large‐scale subduction erosion did not affect the Oligocene‐Miocene Ecuadorian margin (Chapter 2). Differentiation effects causing Y depletion Excluding the Late Miocene Quimsacocha ande‐
sites, all samples characterized by Y <10 ppm rep‐
resent silicic intrusions (SiO2 >64 wt.%). Mineral phases whose stability field is extended in silicic melts, and which are highly compatible for Y comprise amphibole, zircon, and titanite. As noted above, amphibole (as major mineral phase) as well as zircon ±titanite (accessory phases) were observed in all or at least some of the Late Oligocene‐Miocene intrusions. Zircon and titanite are unlikely to saturate in bulk andesitic melts (although they might saturate in local melt pock‐
ets; Hoskin et al. 2000). While amphibole is stable in H2O‐rich mafic melts and may, in concert with garnet, drive Y depletion in mafic‐intermediate melts at deep to mid‐crustal levels (e.g., Richards & Kerrich 2007), its partition coefficient for Y is highly sensitive to melt composition and sharply increases in silicic melts (Bachmann et al. 2005). Fractionation (or restite equilibration) of these minerals thus succeeds in explaining the ob‐
served restriction of strong Y depletion to highly differentiated compositions. To illustrate the petrogenetic significance of am‐
phibole, zircon, and titanite fractionation, some exemplary fractionation trends are displayed in Figure 10. Figure 10a illustrates that, starting at a position typical for a relatively mafic rock compo‐
sition in our dataset (285 ppm Sr; 19 ppm Y), the gross Sr/Y vs. Y distribution trend of the investi‐
gated Late Oligocene‐Miocene intrusions and volcanics could be entirely described by progres‐
sive amphibole fractionation using a Damphibole/melt value for andesitic melts (Rollinson 1993). How‐
ever, using a constant Damphibole/melt value for an andesitic melt requires unrealistically high F val‐
ues (up to 70%) to explain extremely fractionated Sr/Y ratios, and therefore should only apply for moderate Y depletion (c. 10‐20 ppm Y; note that the parental melt composition could also start within this range and subsequently evolve to‐
wards higher Y concentrations by fractionating solely Y‐incompatible minerals such as plagio‐
clase). Starting from a position corresponding to a more silicic rock composition in our dataset (as proxy for an evolved intrusive parental melt; 335 ppm Sr; 9 ppm Y), and using D values for dacitic‐ 141
Figure 9: Trace element and trace element ratio vs. age diagrams. Age‐trace element distribution patterns
show sys‐
tematic increases (Sr, Sm/Dy, Sr/Y) or decreases (Y) in Late Oligocene to Late Miocene arc magmas suggesting progres‐
sively increasing high‐pressure magma differentiation (trends marked by yellow arrows). The latter might be caused by
crustal thickening, or by the downwards migration of crustal hot zones. Extreme Y and HREE depletion at Balsapamba is
a local, anomalous phenomenon, and produces elevated Sr/Y and Sm/Dy ratios in the Early Miocene. Sr depletion for some Junin porphyry intrusions might be caused by shallow crustal plagioclase fractionation or hydrothermal altera‐
tion. Ages estimated from geochronologic constraints summarized in Chapters 2 and 3. NVZ main and frontal arc data from Chiaradia et al. (2009a). See text for discussion. rhyolitic melts (Bachmann et al. 2005), fractional crystallization of small amounts (<10%) of amphi‐
bole, or trace amounts (<1%) of either zircon or titanite, reproduces the observed extreme Y de‐
pletion. Fractionating subequal amounts of pla‐
gioclase and amphibole in a dacitic‐rhyolitic melt results in Y depletion at constant Sr/Y ratios of derivative liquids and might apply for some Junin porphyry intrusions (Fig. 10a), although hydro‐
thermal alteration‐driven Sr depletion is more probable to explain the distribution of Junin por‐
phyries in the Sr/Y vs. Y space. While amphibole, zircon, and titanite fractiona‐
tion in silicic melts affects Sr/Y vs. Y distribution trends in a similar way, their potential individual 142 contributions can be discerned when other trace elements are considered. Fractionating only trace amounts (<1%) of zircon from a dacitic‐rhyolitic melt results in extreme Zr depletion of derivative liquids, which is not observed in our dataset (Fig. 6). Although compatible in both minerals, titanite is characterized by a much lower D value for Sc than amphibole in silicic melts (6 vs. 45: Bach‐
mann et al. 2005; note that pyroxene, although not observed in these intrusions, is also highly compatible for Sc, and its presence in hypothetic parental melts cannot be ruled out, as it is com‐
monly reacted out during late stages of intrusion development to form amphibole or biotite; Bachmann et al. 2007). Therefore, if Sr/Y frac‐
tionation was mainly driven by amphibole frac‐
tionation, it should be accompanied by a propor‐
tionally higher increase in Sr/Sc (as DSc >> DY for amphibole in silicic melts; Bachmann et al. 2005). In contrast, titanite fractionation‐dominated changes in Sr/Y should have a minor effect on Sr/Sc ratios. Figure 10b shows that small amounts (<5%) of combined amphibole‐titanite fractiona‐
tion (in 95:5 proportions) closely reproduce the compositional Sr/Y vs. Sr/Sc range of most silicic intrusions, although a higher proportion of ti‐
tanite with respect to amphibole might be re‐
quired for some intrusions at Balsapamba‐ Telim‐
bela. Several Saraguro Group volcanics (at Cañi‐
capa) are characterized by high Sr/Sc ratios with‐
out a concomitant strong increase in Sr/Y sug‐
gesting Sr/Sc fractionation there was mainly driven by amphibole (or clinopyroxene) fractiona‐
tion. The latter is in agreement with the scarcely reported occurrence of titanite in volcanic rocks (Hoskin et al. 2000). We hence suggest that extreme Y (and HREE) de‐
pletion of Miocene Ecuadorian intrusions was mainly driven by amphibole ±titanite fractiona‐
tion in hydrous silicic melts. As noted above, sig‐
nificant plagioclase fractionation is not observed for these intrusions (mostly REE group 3), such that pronounced magma evolution at shallow crustal levels (<0.4 GPa), where plagioclase is stable closer to the liquidus than amphibole, is unlikely (e.g., Grove et al. 2003). Deep crustal (e.g., at 1.2 GPa) H2O‐rich magma evolution po‐
tentially involves significant garnet fractionation (Alonso‐Perez et al. 2009), which is mostly not observed. Consequently, parental melts to most Miocene intrusions might have dominantly evolved at pressures of c. 0.4‐0.8 GPa where am‐
phibole is the dominant liquidus phase (Alonso‐
Perez et al. 2009). In contrast, Quimsacocha vol‐
canics (and several other samples of REE group 4) share many compositional features with Quater‐
nary NVZ magmas, including a prominent garnet signature in their REE patterns (Fig. 8). Conse‐
quently, Y depletion in these magmas was likely driven by combined amphibole and garnet frac‐
tionation/restite equilibration, possibly at higher pressures than inferred for most Miocene intru‐
sive centers. The petrogenetic significance of Sr/Y ratios for crustal magma evolution Strong Y depletion at broadly constant Sr con‐
tents in the above‐listed Miocene intrusions im‐
plies that Sr/Y ratios of these intrusions are main‐
ly controlled by Y depletion; Sr/Y ratios may in‐
crease towards adakite‐like values (Fig. 7) al‐
though Sr contents are often <400 ppm (the min‐
imum Sr content inferred for most adakite‐like rocks; e.g., Richard & Kerrich 2007). As outlined above, this is mainly due to silicic melt differen‐
tiation by amphibole ±titanite fractiona‐
tion/restite equilibration, although in some cases Y depletion is also influenced by garnet fractiona‐
tion/restite equilibration (REE group 4: Quimsa‐
cocha, Junin porphyries). Adakite‐like geochemical features of modern NVZ magmas in Ecuador have been shown to be mostly the result of crustal magma evolution (e.g., Chiaradia et al. 2009a). In the latter case, high Sr/Y ratios commonly signal high‐pressure magma differentiation, whereas low Sr/Y ratios indicate upper crustal magmatism at low pres‐
sures (e.g., Tulloch & Kimbrough 2003; Bachmann et al. 2005). High Sr/Y ratios (with a threshold value of c. 30‐40) in our dataset are associated with parental melt evolution without significant plagioclase fractionation at shallow crustal levels as inferred from REE distribution patterns (Tab. 3), and correlate positively with Sm/Dy (Fig. 7). In contrast, Sr/Y ratios <30‐40 pool with low Sm/Dy ratios reflecting low‐pressure magma differentia‐
tion. Increasing magma differentiation pressures result in an increasing relative proportion of garnet in an amphibole‐garnet liquidus assemblage for H2O‐rich andesitic melts (Alonso‐Perez et al. 2009). As discussed above, considering variable, melt composition‐dependent partition coeffi‐
cients, extreme Y depletion by amphibole frac‐
tionation only applies to relatively silicic melt compositions, whereas garnet fractionation may produce strong Y depletion in more mafic melts already. Late Miocene (e.g., Quimsacocha) or younger arc magmas evolving by combined am‐
phibole‐garnet fractionation/restite equilibration may therefore develop towards adakite‐like Sr/Y ratios at less silicic compositions, i.e., during ear‐
143
lier differentiation stages than pre‐Late Miocene arc magmas (Fig. 7). ture (Chapter 4) they are shown in separate dia‐
grams in Figure 11. Trace element concentrations and ratios used as proxies for increasing high‐pressure magma dif‐
ferentiation (and further modulated by melt composition and differentiation effects) in Figure 9 show a north‐south division of associated mag‐
matic centers: low‐pressure differentiation is mainly inferred for southern Ecuador (Portovelo‐
Zaruma, Cangrejos, Saraguro), whereas high‐
pressure differentiation mostly applies for north‐
ern‐central Ecuador (Quimsacocha, Apuela‐
Junin). We suggest that this is mainly an effect of the spatio‐temporal distribution of arc magma‐
tism in Ecuador; Mid‐ to Late Miocene arc mag‐
matism in southern Ecuador migrated north‐
wards in response to progressive slab flattening, such that arc magmatic exposures in southern Ecuador are biased towards older arc units and, by inference, an overall thinner crust (Chapter 2). However, regional along‐arc differences might additionally apply, as the potential for tectonic crustal thickening may be higher in northern‐
central Ecuador where the paleo‐continental margin is buttressed against the allochthonous oceanic plateau block (e.g., Jaillard et al. 2005). Southern‐central Ecuadorian Sierra intrusions and volcanics (“east of CPPF” in Fig. 11) in the south‐
wards projection of the NVZ main arc consistently display low radiogenic 87Sr/86Sr values at Sr/Y ra‐
tios >30 (Fig. 11). In contrast, magmas in south‐
ern Ecuador with a significant shallow crustal dif‐
ferentiation step (mainly the Cangrejos‐ Isotopic constraints on shallow vs. deep crustal magma evolution Figure 11 shows plots of several radiogenic iso‐
tope ratios (see Chapter 4 for isotope references) vs. Sr/Y ratio s; Paleogene intrusions and volcan‐
ics in central‐southern Ecuador (Chiaradia et al. 2004) are shown for reference. As discussed above, Sr/Y ratios in our dataset may, to some extent, discriminate between dominant magma evolution at shallow (usually Sr/Y <30) or mid‐ to deep crustal levels (Sr/Y >30) if significant amphi‐
bole ±garnet ±titanite fractionation is involved which is not always the case (e.g., at Gaby). To account for the latter we additionally report the presence of negative Eu anomalies in the legend of Figure 11, indicative of significant shallow crustal plagioclase fractionation. As Miocene in‐
trusions emplaced in the northern‐central Ecua‐
dorian Western Cordillera show significant iso‐
topic differences to central‐southern Ecuadorian Sierra intrusions and volcanics reflecting along‐ and across‐arc differences in basement architec‐
144 Figure 10: Sr/Y vs. Y (A) and Sr/Y vs. Sr/Sc (B) distribu‐
tions of Late Tertiary Ecuadorian granitoids and volcan‐
ics, with theoretical effects of amphibole, titanite, zir‐
con, and plagioclase fractional crystallization (FC) using partition coefficients of Rollinson (1993) for andesitic melts, and Bachmann et al. (2005) for rhyolitic melts. Tick marks (small white diamonds) on andesitic melt amphibole fractional crystallization trend correspond to 10% steps in F values. Fractionation curves for am‐
phibole, titanite, and zircon in silicic melts overlap and are simplified as one in the diagram; whole curve range corresponds to c. 10% amphibole, and <1% ti‐
tanite and zircon FC. See text for discussion.
145
Zaruma intrusive belt and Saraguro Group volcan‐
ics; groups 1±2) extend towards higher 87Sr/86Sr values, analogous to the Paleogene intrusions and volcanics of Chiaradia et al. (2004). Plotting Sr concentrations instead of the Sr/Y ratio (Fig. 11) results in the same distribution pattern where a Sr content of c. 300 ppm corresponds to the threshold value dividing high and low radiogenic maximum 87Sr/86Sr values. This trend is further reflected in a Eu/Eu* vs. 87Sr/86Sr plot (Fig. 11) where lower Eu/Eu* values (shallow crustal magma evolution) again tend to be associated with higher 87Sr/86Sr. Isotope‐based discrimination between shallow and mid‐ to deep crustal magma evolution in southern‐central Ecuador becomes slightly more blurred for 143Nd/144Nd, and significantly more blurred for 207Pb/204Pb (Fig. 11), suggesting that MASH/hot zone‐hosting units in the mid or lower crust have a higher variability in these isotopic ratios. This is mirrored by systematic across‐arc trends in 143Nd/144Nd and 207Pb/204Pb possibly re‐
flecting changes in the mid‐ to deep crustal basement architecture and composition (Chapter 4). Western Cordillera intrusions (“west of CPPF” in Fig. 11) do not display systematic isotopic differ‐
ences for samples with or without negative Eu anomalies and thus, by inference, with or without significant shallow crustal magma evolution. The absence of systematic isotopic offsets might indi‐
cate a vertically relatively homogeneous isotopic composition of the crust below the Western Cor‐
dillera, in agreement with tectonic constraints (e.g., Vallejo et al. 2009). There is a small differ‐
ence, however, between the isotopic composi‐
tions of the Gaby intrusive center and Western Cordillera granitoids situated further north (slightly more radiogenic Sr and less radiogenic Nd at Gaby; Fig. 11) potentially indicating along‐
arc isotopic differences of oceanic plateau units. Overall, the isotopic compositions of Late Tertiary intrusions and volcanics in southern Ecuador clearly reflect polybaric magma evolution at low‐
er to mid‐, and, variably, upper crustal levels. Up‐
per and lower crustal units of the Western Cordil‐
lera seem to be of similar isotopic composition such that isotopes fail to illustrate polybaric evo‐
lution stages there. The visibility of crustal con‐
146 tamination in arc magmas is a function of the compositional and isotopic leverage of potential assimilants; open‐system magma differentiation at the base of the crust may be reflected by only subtle changes in isotopic compositions of deriva‐
tive liquids if the arc position stays stable over geologically long periods of time such that evolv‐
ing arc magmas mainly consume arc intrusive roots (e.g., Davidson et al. 1987). This is reflected by relatively constant 87Sr/86Sr values for intru‐
sions and volcanics with Sr/Y >30 (Fig. 11); only samples with Sr/Y <30 tend to extend to signifi‐
cantly higher 87Sr/86Sr values. Note, however, that primitive isotopic ratios of Late Tertiary Ecuador‐
ian arc magmas do not necessarily reflect assimi‐
lation of arc intrusive roots, but might instead or additionally also indicate assimilation of oceanic plateau units. The latter is clearly visible for Western Cordillera intrusions which show quasi constant 87Sr/86Sr values, but slight to moderate variations in their 143Nd/144Nd and 207Pb/204Pb ra‐
tios, respectively (Fig. 11). Petrogenesis of Late Tertiary arc magmas in Ecuador – summary Building on the constraints and caveats pre‐
sented in the preceding sections, we combined the available petrologic evidence to qualitatively estimate at which crustal levels Late Oligocene to Miocene Ecuadorian arc magmas acquired their geochemical and isotopic characteristics (Tab. 3). Variable parental magma evolutionary paths are reflected by systematic compositional and iso‐
topic differences in Late Tertiary arc magmatic products and allow the distinction of four major groups. These comprise distinct differentiation processes in lower to mid‐crustal hot zones, as well as variable shallow crustal overprinting, and are schematically sketched in Figure 12. (1) Most intrusions and volcanics characterized by a significant shallow crustal magma evolution stage occur in southern‐central Ecuador (Sara‐
guro Group volcanics and the Cangrejos‐Zaruma intrusive belt) and have Early to Mid‐Miocene ages (Chapter 2). These lithologies show minor‐
moderate negative Eu anomalies and variable, relatively high Y and HREE contents (11‐25 ppm Y, Tab. 3; note, though, that most of these values still qualify as adakite‐like), variable Sr contents Figure 11: Radiogenic Sr, Nd, and Pb isotopic composition vs. Sr/Y (plus Sr and Eu/Eu*) distribution of Late Tertiary Ec‐
uadorian granitoids and volcanics. Plots are divided according to different upper crustal compositions (continental vs. oceanic, i.e., east vs. west of the Calacali‐Pallatanga‐Pujili fault zone, respectively) to better resolve isotopic variations. As discussed in the text, the Sr/Y ratio serves as a proxy to discriminate lower vs. upper crust‐dominated magma evolu‐
tion (threshold value for our dataset = c. 30). Paleogene intrusions and volcanics of Chiaradia et al. (2004) are shown for reference; they consistently overlap with distribution trends of Late Tertiary magmas where significant shallow
crustal magma evolution is inferred suggesting a broadly similar petrogenesis. See text for discussion. 147
(228‐459 ppm), and Sr/Y ratios in the range of 13‐
32. Their composition is consistent with H2O‐
undersaturated parental melt evolution at basal to mid‐crustal levels of a relatively thin to mod‐
erately thick arc crust in the Late Oligocene to Early Miocene where pyroxene constitutes the dominant liquidus phase (e.g., Müntener et al. 2001), followed by shallow crustal‐dominated magma evolution involving significant plagioclase fractionation. While their REE patterns mostly do not show any prominent MREE‐HREE fractiona‐
tion, limited magma evolution under influence of a Y‐fractionating mineral is required to explain the variable degrees of Y depletion prior to, or concomitant with plagioclase fractionation. (2) A second group of intrusions and volcanics in southern Ecuador overlaps with the previous group in time and space (often forming part of the same integrated intrusive center), and their main difference to the previous group is the ab‐
sence of negative Eu anomalies in their REE dis‐
tribution patterns, suggesting the lack of signifi‐
cant shallow crustal plagioclase fractionation, or, alternatively, indicating highly oxidized melts where plagioclase fractionation did not result in negative Eu anomalies. The lower to mid‐crustal petrogenesis of these magmas is similar to the previous group. At roughly constant Sr and slightly lower Y contents, Sr/Y ratios of this group are slightly higher than in the previous group (Tab. 3). The integrated vertical magma evolution of these groups (1 and 2) is clearly recognizable in 87
Sr/86Sr, and to a lesser extent in 143Nd/144Nd iso‐
topic compositions (Fig. 11), and supports the principal notion of Chiaradia et al. (2009a) that isotopic differences between Tertiary and Qua‐
ternary NVZ magmas (in northern‐central Ecua‐
dor) are partly controlled by the crustal depth of magma evolution. Additionally, our regionally more representative dataset of Late Tertiary magmas allows us to refine this conclusion in two points: (1) the Miocene magma evolution in southern‐central Ecuador does not always in‐
clude significant shallow crustal magma differen‐
tiation, although it is common; and (2) the iso‐
topic composition of the mid‐ to upper (and lower?) crust might be regionally heterogeneous and thus can further modulate the observed iso‐
topic distribution patterns. The latter is sug‐
148 gested by the isotopic compositions of some Late Tertiary magmas (mainly the Cangrejos‐Zaruma intrusive belt), which are significantly more ra‐
diogenic than any known isotopic composition of present‐day NVZ magmas (compare Chapter 4), in agreement with the inferred continental basement domains of the southern Ecuadorian El Oro block (Litherland et al. 1994). (3) A third group of Miocene intrusions, mainly hosted by the Western Cordillera and its western foothills in northern‐central Ecuador, is charac‐
terized by strong Y and HREE depletion (mostly with concave‐upwards HREE patterns) and con‐
sistently displays either minor‐moderate positive Eu anomalies, or smooth Sm‐Eu‐Gd transitions, indicative of amphibole ±titanite fractiona‐
tion/restite equilibration without significant pla‐
gioclase or garnet fractionation. Strontium con‐
tents are in the 300‐600 ppm range. Depending on the degree of Y depletion, Sr/Y ratios range from 31 up to 282 (Tab. 3) and thus qualify as adakite‐like (Fig. 7), although Sr contents are mostly below the threshold for adakite‐like com‐
positions (<400 ppm). Strong Y depletion by am‐
phibole ±titanite fractionation occurs in silicic melts where titanite may be stable (Hoskin et al. 2000) and DY values of amphibole increase sig‐
nificantly (e. g., Bachmann et al. 2005). Crystallization experiments on hydrous andesitic melts show that amphibole is the dominant liq‐
uidus phase at moderate crustal depth (e.g., at 0.8 GPa corresponding to c. 25 km depth; Alonso‐
Perez et al. 2009). A change from pyroxene‐ to amphibole‐dominated magma evolution at depth (groups 1 and 2 vs. 3) is commonly associated with increasing hot zone depth (e.g., Kay & Mpo‐
dozis 2002). Additionally, amphibole becomes destabilized towards lower melt H2O contents at a given temperature and pressure (e.g., Müntener et al. 2001; Alonso‐Perez et al. 2009). Consequently, in addition to reflecting an in‐
crease in hot zone depth (by crustal thickening or by downward shifting of the hot zone position(s) from mid‐ towards deep crustal levels), the change from pyroxene‐ to amphibole‐dominated magma differentiation between the two arc seg‐
ments might also be related to overall higher melt H2O contents. Whereas crustal thickening can be considered as a regional‐scale effect, vari‐
ations in melt H2O contents could also apply at a Figure 12: Schematic illustration of transcrustal petrogenesis of Late Oligocene to Mid‐Miocene (left) and Late Miocene to Quaternary (right) Ecuadorian arc magmas. Arc magmas in both periods are generally processed in lower to mid‐
crustal MASH/hot zones (orange bars); crustal thickness, hot zone depth, and melt water contents control the lower crustal petrogenesis in each case and may variably include amphibole ± garnet (± titanite) fractionation or restite equilibration, respectively. Involvement of a subsequent shallow crustal magma evolution step (characterized by pla‐
gioclase fractionation) may depend on multiple regional‐local factors such as magma supply rate and crustal heat anomalies (during batholith construction), or stress field. Adakite‐like arc magmas form by high‐pressure crustal differ‐
entiation of H2O‐rich magmas without subsequent significant low‐pressure differentiation. Additional variations may be caused by compositional changes of mantle wedge‐derived melts invading lower crustal hot zones, in particular since the Late Miocene. The indicated Late Oligocene to Mid‐Miocene crustal thickness is only a rough estimate; the present‐
day 50‐70 km thickness is constrained by seismic studies (Guillier et al. 2001). See text for further discussion. rather local scale (e.g., Rodriguez et al. 2007). This might partly explain the contrasting lower to mid‐crustal magmatic evolution at Gaby and Balsapamba, where only the latter involves sig‐
nificant amphibole fractionation or restite equili‐
bration. Both intrusive systems share a number of similar features which include age (c. 20 Ma; Chapter 2), a frontal arc position hosted by the same basement lithology (oceanic plateau), and moderately‐highly differentiated compositions (SiO2 c. 61‐68 wt.%) where plagioclase and horn‐
blende constitute the main phenocrysts; these broad similarities suggest that local (in addition to regional) factors can exert a major control on the lower to mid‐crustal petrogenesis for paren‐
tal melts of a given magmatic center. (4) A last group comprises Late Miocene intru‐
sions and volcanics (Junin porphyries and Quim‐
sacocha) which are characterized by strong HREE depletion and negative HREE slopes indicating combined amphibole and garnet fractiona‐
tion/restite equilibration during their petroge‐
netic evolution (e.g., H2O‐rich melts processed in a mid‐crustal hot zone at 0.8 GPa; Alonso‐Perez et al. 2009). Strontium contents are high at Quimsacocha (448‐858 ppm), but low at Junin (<336 ppm), the latter probably influenced by hydrothermal alteration. Negative Eu anomalies 149
are not present suggesting the absence of signifi‐
cant shallow crustal plagioclase fractionation sub‐
sequent to MASH/hot zone processing of Late Miocene arc magmas. The petrogenetic evolution of Late Miocene arc magmas (group 4) resembles several present‐day NVZ volcanic centers where Chiaradia et al. (2009a) propose that bulk arc compression caus‐
es their parental melts to evolve at deep crustal levels of a thick crust; a similar mechanism might apply to Late Miocene arc magmas in northern‐
central Ecuador. However, while Chiaradia et al. (2009a) propose collision of the Carnegie Ridge with the Ecuadorian margin as the cause for bulk margin compression, the latter can only apply for arc magmas <8 Ma, the maximum age of the in‐
ception of Carnegie Ridge collision (Chapter 3). Furthermore, initial ridge collision took place in northern Ecuador; the ridge axis then progres‐
sively swept southwards along the margin (Chap‐
ter 3). Carnegie Ridge subduction thus cannot have affected the far‐field stress regime of the central Ecuadorian arc in the Late Miocene such that alternative causes of compression are neces‐
sary to explain, for example, pronounced deep crustal magma evolution at the 7 Ma Quimsaco‐
cha volcanic center in central Ecuador. The latter could be related to local compression in a region‐
ally restraining bend structural setting associated with oblique plate convergence at the Ecuadorian margin (Chapter 2). Alternatively, compression might not always be required to cause pro‐
nounced magma evolution at mid‐ to deep crustal levels. Figure 13: Sr/Y vs. Y distribution of Late Tertiary por‐
phyry intrusions and spatially associated host intrusive
units in Ecuador. Except for the Cangrejos porphyry
intrusions, all porphyries lack petrogenetic evidence
for shallow crustal plagioclase fractionation. Amphi‐
bole ±titanite ±garnet fractionation/restite equilibra‐
tion at mid‐crustal levels drives some porphyry paren‐
tal melts towards adakite‐like compositions. Mid‐
crustal magma evolution towards the end of magmatic cycles of batholith construction seems to represent a
metallogenically favorable environment whereas intense shallow crustal magmatism during batholith peak assembly
may be less favorable to form and preserve porphyry‐related mineralization. Overall, the occurrence of adakite‐like features in porphyry intrusions follows the regional compositional characteristics of arc magmatism at a given time. If
present, trace element compositional differences between porphyries and associated host intrusions are associated
with regional arc magmatic trends through time. Intrusive ages from Chapters 2 and 3, and references therein. See text for further discussion. 150 Significance of adakite‐like fea‐
tures for Late Oligocene to Late Miocene porphyry‐related min‐
eralization in Ecuador A spatial association of adakite‐like magma com‐
positions and porphyry‐related mineralization has been observed in a number of studies of circum‐
Pacific subduction‐related ore deposits (e.g., Thiéblemont et al. 1997; see Richards & Kerrich 2007 for a comprehensive summary and an evaluation of metallogenic implications). Petro‐
genetic controls for the development of adakite‐
like features in Late Tertiary Ecuadorian arc magmas have been discussed above and com‐
prise a combination of regional‐scale (hot zone depth, total crustal thickness) to local‐scale (melt H2O contents, differentiation effects in silicic magmas including accessory phases) factors. As noted above, adakite‐like features (Sr/Y ratios >30‐40) in our dataset indicate the absence of extensive shallow crustal plagioclase fractiona‐
tion, and the fractionation (or restite equilibra‐
tion) of amphibole ±garnet ±titanite at higher pressure or at relatively high melt H2O contents. As such, the above‐mentioned observation of adakite‐like magmatism associated with por‐
phyry‐related mineralization would imply that pronounced deep to mid‐crustal H2O‐rich melt equilibration with mainly amphibole, and the ab‐
sence of a pronounced shallow crustal magma evolution step generated favorable conditions for subsequent porphyry‐related mineralization from fluids exsolved from these melts at somewhat shallower depth. The latter has been demon‐
strated at the Late Miocene‐Pliocene Tampakan porphyry Cu/epithermal high sulfidation Cu‐Au ore deposit district (Mindanao, Philippines) where parental melt evolution involves pro‐
nounced amphibole fractionation at moderate crustal depth (0.5‐0.6 GPa) whereupon the melt becomes enriched in H2O (and Cl); the latter re‐
sults in earlier, i.e., higher pressure volatile satu‐
ration of subsequently ascending melt batches where exsolved fluids are highly saline and Cu‐
rich (Rohrlach & Loucks 2005). To test the significance of this association for Late Tertiary Ecuadorian porphyry systems, we plotted the compositions of porphyry intrusions and as‐
sociated precursor intrusions in Sr/Y vs. Y dia‐
grams for all five major porphyry systems (Junin, Balsapamba‐Telimbela, Chaucha, Gaby, Cangre‐
jos) investigated in this study (Fig. 13). Porphyry stocks at Junin, Balsapamba‐Telimbela, and Chaucha are characterized by adakite‐like signa‐
tures, whereas Gaby and Cangrejos are not (Fig. 13). At the individual deposit scale, the occur‐
rence of adakite‐like magmatism seems to be broadly temporally controlled (see Chapters 2 and 3 for age references): there is a shift towards adakite‐like magma compositions at Junin (from >15 Ma to 13‐6 Ma; note that Sr contents of Junin porphyries might have been lowered by hydrothermal alteration such that these intru‐
sions would originally plot at higher Sr/Y ratios), at Telimbela (from c. 26 Ma to 20 Ma), and at Chaucha (from 15 to 10 Ma). In contrast, no sys‐
tematic shifts are recorded at Gaby and Cangre‐
jos (always low Sr/Y) and Balsapamba (always high Sr/Y) where the intrusive evolution spans a relatively short time range of <2 m.y. (note that Balsapamba batholith growth comprises a signifi‐
cantly longer time span, but geochemical data for older intrusive pulses are not available). The distribution of adakite‐like features of Late Tertiary porphyry systems along the Ecuadorian margin demonstrates that porphyry‐related min‐
eralization is not exclusively associated with a specific geochemical signature or a specific path of crustal magma evolution in Ecuador; any arc magma may potentially form porphyry‐related mineralization in a favorable tectonomagmatic setting. The spatiotemporal distribution of ada‐
kite‐like features in Ecuadorian porphyry intru‐
sions largely reflects regional temporal trends in arc magma geochemistry. In addition, the con‐
trasting chemical signatures (adakite‐like vs. non‐
adakitic) do not show any first‐order basement control, as intrusions and porphyry systems of both groups are mainly hosted by oceanic pla‐
teau units or Tertiary arc volcanics and granitoids. Only the Cangrejos igneous complex and the as‐
sociated porphyry system show isotopic evidence of significant magma contamination by continen‐
tal crust (Chapter 4). There are a number of broadly applicable systematic distribution criteria between the various porphyry systems: 
If Balsapamba‐Telimbela is excluded, Sr/Y ratios tend to become broadly higher with 151
decreasing intrusive age on a regional scale (influencing the age distribution systemat‐
ics at the deposit scale as noted above). 
Porphyry systems in Figure 13 are arranged from north to south; bulk Sr/Y ratios of porphyry intrusions tend to become higher towards the north. If the anomalous occur‐
rence of adakite‐like features at Balsa‐
pamba‐Telimbela is excluded, however, the north‐south trend corresponds to a pro‐
gressively increasing age distribution, con‐
sistent with the previous point. 
Except for Cangrejos, porphyry intrusions lack negative Eu anomalies indicating the absence of significant shallow crustal pla‐
gioclase fractionation. 
All porphyry systems characterized by high Sr/Y ratios (Junin, Balsapamba, Telimbela, Chaucha) represent Cu‐Mo porphyry de‐
posits (Prodeminca 2000), whereas Au‐Cu porphyry systems (Gaby, Cangrejos) show low Sr/Y ratios. The geochemical composition of most porphyry intrusions points to the absence of significant shallow crustal magma evolution of their parental melts; where porphyry intrusions are associated with larger intrusive complexes, porphyry em‐
placement is late with respect to batholith con‐
struction (Chapters 2, 3). Peaking of the latter is commonly associated with high magma supply rates and might have allowed establishment of large shallow crustal magma chambers involving major plagioclase fractionation (e.g., Bachmann et al. 2005, 2007). Dwindling magma supply rates might eventually cause a downwards collapse of the focus of magmatism towards greater depth (pressure) where amphibole is stable closer to the liquidus than plagioclase (e.g., Grove et al. 2003); magma subsequently intruding into the shallow crust cools rapidly below its solidus with‐
out significant further differentiation (e.g., Annen et al. 2006). As discussed by Rohrlach & Loucks (2005), progressive melt volatile enrichment by magma evolution (and replenishment) at moder‐
ate pressure favorably influences fluid exsolution kinetics (and pressure‐dependent melt‐fluid par‐
titioning of Cl as a major Cu complexing agent) of subsequently ascending melt batches with re‐
spect to porphyry‐related mineralization. The 152 underrepresentation of long‐lived shallow crustal magmatic systems directly associated with por‐
phyry intrusions might relate to (1) lower H2O solubilities of melts stalled at lower pressures such that less parental melt preconditioning by volatile enrichment takes place; for a given melt volume, the overall potential volume of exsolved fluid focused in space and time, and thus the size of the porphyry‐related hydrothermal system, would then decrease; (2) volatile loss to the sur‐
face by volcanism, and fluid dissipation instead of focused flow; (3) destruction of shallow crustal mineralization by later intrusive pulses. As such, adakite‐like compositions of porphyry intrusions might signal favorable tectonomagmatic precon‐
ditioning of porphyry parental melts for subse‐
quent intrusion‐related mineralization. The observed mutual exclusivity of Cu‐Mo and Au‐Cu porphyry systems with or without adakite‐
like magma compositions, respectively, may be an apparent one, as adakite‐like features are as‐
sociated with Au‐Cu or Cu‐Au por‐
phyry/epithermal mineralization elsewhere (e.g., Rohrlach & Loucks 2005; Chiaradia et al., 2009b). However, it might indicate that deep to mid‐
crustal magma evolution and the inferred con‐
comitant volatile enrichment (and its bearing on the relative timing of volatile saturation and fluid exsolution depth) of porphyry parental melts may be particularly important for the Cu‐Mo budget of a given porphyry system, whereas additional factors influence its Au budget. The significance of magma chemistry for the total tonnage of a given ore deposit in Ecuador cannot be accurately evaluated as the Late Tertiary porphyry systems are variably deeply eroded such that their current tonnage does not necessarily reflect the initial deposit size (Prodeminca 2000). Conclusions The overall spatio‐temporal distribution of ada‐
kite‐like features in Late Tertiary Ecuadorian arc magmas is semi‐systematic; magmatic centers characterized by (partly) adakite‐like magmatism are mainly hosted by the Western Cordillera in northern‐central Ecuador and comprise Balsa‐
pamba (c. 21 Ma), Apuela‐Junin (13‐6 Ma), Chau‐
cha (c. 10 Ma), and Quimsacocha (7 Ma). Adakite‐
like features (high Sr/Y) of Late Tertiary Ecuador‐
ian arc magmas are mainly due to strong Y and heavy REE depletion of their parental melts at broadly constant Sr contents, and are related to fractionation/restite equilibration effects of am‐
phibole, garnet, and titanite. In the Early to Mid‐Miocene, amphibole (± acces‐
sory titanite) is the most important mineral phase for controlling Y and HREE depletion in silicic arc magmas and thus their evolution towards ada‐
kite‐like features in Ecuador, either by fractiona‐
tion, or equilibration with an amphibole‐bearing restite. The onset of Y and heavy REE depletion by garnet fractionation/restite equilibration (in concert with amphibole) seems to be restricted to the Late Miocene and continues to the present day. While strong Y depletion by amphibole frac‐
tionation/restite equilibration is particularly effi‐
cient in silicic melts, Y depletion by garnet frac‐
tionation/restite equilibration is also efficient in mafic melts and allows Late Miocene and younger arc magmas to acquire adakite‐like com‐
positions already at somewhat earlier differentia‐
tion stages than Late Oligocene to Mid‐Miocene magmas. Shallow crustal plagioclase fractionation affects some, but not all Late Tertiary arc mag‐
mas in southern Ecuador; it is of minor petroge‐
netic significance for Miocene intrusions of the Western Cordillera in northern‐central Ecuador. A preferential association of adakite‐like features with a specific basement lithology cannot be ob‐
served. Systematic trace element variations (Sr, Y, REE) through time are indicative of progressively in‐
creasing high‐pressure crustal magma differentia‐
tion. While increasing crustal thickness favorably influences the occurrence of adakite‐like fea‐
tures, the latter are further modulated by a set of parameters which dynamically control mineral stabilities and mineral‐melt partitioning coeffi‐
cients. These include magma evolution depth (pressure) in a given crustal column and melt composition, the latter comprising the degree of differentiation and melt H2O contents; these ad‐
ditional controls may operate at a regional or lo‐
cal scale. For porphyry‐related ore deposits where multi‐
m.y. precursor magmatism occurs at the same site, and porphyry emplacement represents the last major intrusive event, we note a tendency of porphyry parental melts to evolve towards ada‐
kite‐like compositions (e.g., at Junin and Chau‐
cha) indicative of downward migration of the fo‐
cus of crustal magma evolution towards greater depth and/or increasing melt H2O contents. While this may reflect favorable tectono‐
magmatic preconditioning of porphyry parental melts, it is important to note that these composi‐
tional changes are of a regional arc magmatic scale, and broadly controlled by the relative age difference between porphyry and host intrusions. Systematic compositional changes between por‐
phyry and precursor intrusions are not recorded if the time difference between their respective emplacement events is small (<2 m.y.). In the lat‐
ter case, magmas may be both of adakite‐like (Balsapamba) or non‐adakitic affinity (Gaby, Can‐
grejos). The fact that porphyry‐related minerali‐
zation in Ecuador spans the whole Late Oligocene to Late Miocene (24‐6 Ma) over a large latitudinal range (c. 0° to 3°30’S) supports the notion that any arc magma of a sufficient volume has the po‐
tential to form porphyry‐related mineralization. Porphyry‐related ore deposits in Ecuador may comprise multiple intrusive phases (Chapters 2, 3), and more detailed studies are necessary to better resolve individual deposit‐scale geochemi‐
cal trends and their significance for mineraliza‐
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related hydrothermal systems Porphyry‐related ore deposits typically comprise multiple intrusive bodies of variable geometry such as cylindrical stocks (commonly <1 km di‐
ameter, but several km in length), dikes, or domes (Seedorff et al. 2005). The intrusive bodies act as focal points for the flow of metal‐bearing magmatic fluids exsolved from a magma chamber situated at greater depth (Sillitoe 1973). These fluids generate hydrothermal alteration of vari‐
able intensity both inside the porphyry body and in a significantly larger aureole surrounding the intrusion. Fluid evolution by fluid ascent, cooling, mixing with external (mainly meteoric) fluids and fluid‐wall rock equilibration dictates its chemical character and the resulting type of hydrothermal alteration (Seedorff et al. 2005). Typical altera‐
tion zones associated with porphyry Cu deposits comprise sodic‐calcic and potassic alteration at depth, giving way to phyllic and argillic alteration at higher or more distant levels, and grading into propylitic alteration in the peripheral parts of the porphyry system. Fluid phase separation into co‐
existing vapor and brine phases additionally re‐
sults in concomitant advanced argillic alteration by the ascending vapor in the uppermost part of the porphyry system. Similar alteration zones ex‐
ist in epithermal deposits. A collapse of the fluid flow pattern driven by decreasing temperatures and fluid supply rates can promote overprinting of the different alteration types. For a compre‐
hensive summary on mineral assemblages asso‐
ciated with these different types of alteration see Seedorff et al. (2005). Hydrothermal alteration produces metasomatic changes of magmatic rock compositions (e.g., Ulrich & Heinrich 2002). As hydrothermal altera‐
tion is inevitable in a porphyry environment, sampling for our study focused on “least altered samples” which are supposed to be closest to original magmatic compositions. In this context, this term has a relative meaning, depending on the local alteration style and intensity, such that alteration‐induced compositional changes in samples labeled as “least‐altered” cannot neces‐
sarily be predicted with confidence (Giffkins et al. 2005). Therefore, for a study of geochemical fea‐
tures of rocks subject to hydrothermal alteration such as ours it is useful to first evaluate element mobility during hydrothermal alteration stages and the resulting changes in rock geochemistry to define the degree to which elements can be used as reliable tracers for petrogenetic processes. While it is generally accepted that the group of large ion lithophile elements (LILE) and, to a less‐
er extent, light rare earth elements (LREE), are readily transportable by fluids, whereas high field strength elements (HFSE) and mid to heavy rare earth elements (M/HREEE) tend to be fluid‐
immobile, detailed geochemical studies of al‐
tered rocks show that HFSE and M/HREE can also be mobilized to significant extents under certain conditions (e.g., Lesher et al. 1985; Fulignati et al. 1999; Polat & Hofmann 2003; Verma et al. 2005; Shikazono et al. 2008). We tested the compositional scatter caused by hydrothermal alteration for a number of litholo‐
gies where multiple samples of the same lithol‐
ogy are available (Apuela‐Cuellaje, Balsapamba and Gaby intrusions; Saraguro volcanics at Por‐
tovelo) by calculating isocons based on least‐
altered samples of each lithology (Grant 1986). At Gaby, samples comprise two different porphyry intrusions which are compositionally indistin‐
guishable, such that they may be treated as one for evaluating alteration‐induced compositional changes. Isocon slopes are calculated from Al2O3 contents of least‐altered versus altered samples, as this oxide consistently shows constant con‐
tents irrespective of style and intensity of altera‐
tion, attesting to its immobile nature. We limit isocon calculations to major element oxides and a number of trace elements serving as important tracers for certain petrogenetic processes. Some mineralogical features of variably altered samples are exemplified in Fig. 2; isocon plots are shown in Fig. 3, and a complete overview table where the relative, isocon‐based compositional changes are quantified can be found in Table A2. Compositional scatter between samples of a given lithology only becomes statistically signifi‐
cant at relative differences exceeding 10% corre‐
sponding to the analytical precision of trace ele‐
ment analysis. To account for the intense altera‐
tion typically associated with porphyry systems, we double this value and define a 20% relative 157
scatter in element contents of the same lithology as acceptable, while regarding elements consis‐
tently showing higher scatter as unreliable to characterize magmatic processes. Using this clas‐
sification scheme, we find the element behavior patterns listed in Table 2 which we assume as representative for our whole dataset. Elements of the LILE (Cs, Rb, Ba, K) and REE (especially LREE) groups show alteration‐induced scatter beyond acceptable means and are thus inappro‐
priate to constrain petrogenetic processes. This inhibits usage of crustal contamination‐sensitive incompatible element ratios such as Rb/K, as well as La/Yb as an overall indicator of REE fractiona‐
tion. Strontium, though part of the LILE group, shows acceptable scatter and its concentrations are therefore considered as petrogenetically signifi‐
cant. The relative immobility of Sr in most sam‐
ples is likely an effect of avoiding intense feldspar phenocryst‐destructive alteration types during sampling, although potential alteration of small feldspar crystals in the submicroscopic porphyry matrix cannot be evaluated. Biotite constitutes a dominant mineral in potassic alteration assem‐
blages in Ecuador; substitution of biotite‐hosted interstitial‐site K with Rb, Cs and Ba, but not Sr (Deer et al. 1992) is consistent with the inferred high mobility of these elements compared to the relative immobility of Sr in our samples. Chemi‐
cally correlated behavior of MREE and HREE in altered rocks suggests that element ratios based on these groups (e.g., Sm/Dy, Dy/Yb) are rela‐
tively non‐susceptible to alteration. Several major elements and the HFSE show only minor to mod‐
erate variations for rocks affected by the various alteration types listed in Table 2, such that their concentrations in hydrothermally altered rocks largely reflect magmatic processes. References Deer, W. A., Howie, R. A., Zussman, J. (1993): An in‐
troduction to the rock‐forming minerals. Pera‐
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earth element (REE) behaviour in the alteration facies of the active magmatic‐hydrothermal system of Vul‐
cano (Aeolian Islands, Italy). Journal of Volcanology and Geothermal Research 88; 325‐342. 158 Giffkins, C., Herrmann, W., Large, R. (2005): Altered volcanic rocks: a guide to description and interpreta‐
tion. Centre for Ore Deposit Research Publication, University of Tasmania; 275 p. Grant, J. A. (1986): The isocon diagram – a simple solu‐
tion to Gresens' equation for metasomatic alteration. Economic Geology 81; 1976‐1982. Lesher, C. M., Goodwin, A. M., Campbell, I. H., Gorton, M. P. (1986): Trace‐element geochemistry of ore‐
associated and barren, felsic metavolcanic rocks in the Superior Province, Canada. Canadian Journal of Earth Sciences 23; 222‐237. Polat, A. & Hofmann, A. W. (2003): Alteration and geochemical patterns in the 3.7‐3.8 Ga Isua green‐
stone belt, West Greenland. Precambrian Research 126; 197‐218. Seedorff, E., and seven others (2005): Porphyry depos‐
its: characteristics and origin of hypogene features. Economic Geology 100th Anniversary Volume; 251‐
298. Shikazono, N. and six others (2008): Geochemical be‐
haviour of rare earth elements in hydrothermally al‐
tered rocks of the Kuroko mining area, Japan. Journal of Geochemical Exploration 98; 65‐79. Sillitoe, R. H. (1973): The tops and bottoms of por‐
phyry copper deposits. Economic Geology 68; 799‐
815. Ulrich, T. & Heinrich, C. A. (2002): Geology and altera‐
tion geochemistry of the porphyry Cu‐Au deposit at Bajo de la Alumbrera, Argentina. Economic Geology 97; 1865‐1888. Verma, S. P., Torres‐Alvarado, I. S., Satir, M., Dobson, P. F. (2005): Hydrothermal alteration effects in geo‐
chemistry and Sr, Nd, Pb, and O isotopes of magmas from the Los Azufres geothermal field (Mexico): a sta‐
tistical approach. Geochemical Journal 39; 141‐163. Appendix II – Data tables Table A1 (next 15 pages): Concentrations of major and trace elements of Late Tertiary magmatic centers. Analyzed elements not described/discussed in Chapter 5 (mostly metals and S) are listed at the end of the table. Analyses of several basement units, xenoliths, and the Late Cretaceous Curiplaya porphyry in‐
trusions which were carried out in the frame of this PhD project are shown for comparison, but are not discussed in the text. Mineral abbreviations used throughout Table A1: qtz ‐ quartz, plag ‐ plagioclase, fsp ‐ feldspar, hbl ‐ horn‐
blende. Other abbreviations: CCOP ‐ Caribbean‐Colombian oceanic plateau; XRF analysis: Rho. ‐ Rhodes pro‐
gram; tr. ‐ standard trace element program. All analyses performed at the University of Lausanne. Tab. A2: Relative changes in concentration compared to least‐altered reference sample concentration calculated from Grant (1986); isocon constructed assuming constant mass of Al2O3 159
Table A1 (continued)
Sample
Method
Location
E05129
E06209
E07034
Apuela batholith at Junin
E06200
E06202
E06205A
Apuela batholith at Cuellaje
E06206A
E06206B
Lithology
granodiorite
granodiorite
qtz-diorite
granodiorite
granodiorite
granodiorite
granodiorite
granodiorite
69.52
0.41
14.87
3.82
0.02
1.58
1.67
3.05
3.34
0.10
1.21
99.6
68.27
0.41
15.57
4.37
0.04
1.70
3.07
3.52
2.53
0.10
0.81
100.4
246
<2
289
4
84
18
87
22
176
28
14
<3
12
29
15
4
65.69
0.44
16.55
3.73
0.05
1.89
4.37
4.44
1.50
0.14
0.51
99.3
2.5
0.14
71
7.5
2.0
1.2
25
479
577
2
0.78
1.5
9.1
18
2.2
9.7
1.9
0.53
1.5
0.21
1.31
0.24
0.69
0.09
0.62
0.10
9
84
28
9.8
18
8
38
18
<3
0.36
<0.06
0.32
1.1
0.16
3.0
66.67
0.43
16.10
3.81
0.06
1.85
4.14
4.49
1.55
0.14
0.32
99.6
2.6
0.15
82
7.1
2.2
1.0
28
654
555
3
0.84
2.3
11
20
2.4
10
2.0
0.60
2.0
0.25
1.2
0.22
0.74
0.10
0.82
0.09
9
81
40
9.5
16
63
54
19
4
0.31
0.08
0.84
0.17
68.12
0.39
15.43
3.39
0.05
1.62
3.60
4.42
1.69
0.13
0.30
99.2
14
66.22
0.41
16.58
3.60
0.04
1.69
3.99
4.42
1.46
0.14
1.19
99.7
2.4
0.16
79
5.9
2.1
1.8
29
1406
563
<2
1.0
1.7
8.5
17
2.0
9.0
1.9
0.55
1.5
0.17
0.91
0.19
0.47
0.06
0.52
0.09
8
81
27
8.5
13
36
38
18
<3
0.47
<0.10
66.49
0.40
16.21
3.63
0.06
1.86
4.26
4.50
1.71
0.13
0.38
99.6
14
64.38
0.54
16.31
5.97
0.04
2.31
3.46
3.10
1.97
0.13
0.98
99.2
2.6
0.22
104
14
3.0
3.7
50
211
331
3
1.7
3.9
13
25
2.9
12
2.6
0.61
2.2
0.36
2.1
0.39
1.4
0.22
1.3
0.23
15
126
29
12
11
191
46
18
4
2.1
0.27
1.9
1.0
288
329
23
104
8
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
160 XRF
XRF
XRF
XRF
XRF
XRF
XRF
XRF
XRF
XRF
XRF
XRF
ICP-MS
ICP-MS
ICP-MS
XRF-Rho.
ICP-MS
ICP-MS
ICP-MS
ICP-MS
XRF-Rho.
XRF-Rho.
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
XRF-Rho.
XRF-Rho.
ICP-MS
ICP-MS
XRF-tr.
XRF-Rho.
XRF-Rho.
XRF-tr.
ICP-MS
ICP-MS
ICP-MS
ICP-MS
ICP-MS
XRF-tr.
<3
12
6.9
549
4
70
26
2
47
19
<3
107
7.4
489
4
69
30
10
52
18
4
Table A1 (continued)
Sample
E05120
E05127
E06183
Location
Junin porphyry intrusions
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
E06211
E07032
E06186
E06197
E07035
E06127 E06130
Balsapamba batholith
fsp-qtz
porpyhry
fsp-qtz
porphyry
porphyritic
microdiorite
tonalite
granodioritic
porphyry
granodioritic
porphyry
granodioritic
porphyry
granodioritic
porphyry
granodioritic
porphyry
69.30
0.23
16.26
1.32
0.00
0.51
0.17
3.39
5.77
0.09
1.24
98.3
0.7
0.06
50
2.4
1.4
0.4
46
620
129
<2
0.46
0.34
1.8
3.9
0.5
2.2
0.34
0.13
0.36
0.04
0.35
0.07
0.16
0.04
0.18
0.02
2
36
<2
1.3
1.3
3536
20
18
3
64
0.43
70.45
0.24
15.94
1.74
0.02
0.76
0.72
5.69
2.91
0.10
0.95
99.5
1.7
0.13
68
3.9
1.9
1.4
40
636
336
<2
0.70
0.82
5.9
12
1.6
7.1
1.4
0.35
1.4
0.15
0.81
0.19
0.41
0.07
0.43
0.06
5
46
7
3.1
4.4
357
26
17
<3
6.3
0.38
69.13
0.22
16.70
2.35
0.01
0.82
0.21
4.00
3.39
0.02
2.33
99.2
68.66
0.33
16.33
2.93
0.02
1.15
0.39
4.67
3.99
0.12
1.39
100.0
1.8
0.10
74
4.4
1.8
0.9
51
624
218
<2
0.88
0.98
8.1
17
2.0
8.7
1.9
0.63
1.2
0.16
0.71
0.14
0.36
0.04
0.24
0.05
5
56
5
6.5
6.9
617
86
18
6
2.0
0.23
70.84
0.24
16.01
2.09
0.02
0.72
0.45
5.61
2.74
0.09
1.32
100.1
1.8
0.12
71
4.1
2.0
0.7
38
543
298
3
0.48
0.95
7.2
15
1.8
8.0
1.3
0.49
1.2
0.14
0.82
0.13
0.36
0.04
0.48
0.06
5
43
6
6.5
5.3
291
29
18
18
1.2
0.22
70.22
0.20
17.02
1.92
0.00
0.65
0.10
2.80
5.15
0.06
1.80
99.9
0.7
0.04
53
2.4
1.5
0.3
55
376
87
<2
0.45
0.55
2.6
4.9
0.6
2.3
0.51
0.13
0.33
0.05
0.26
0.06
0.19
0.04
0.24
0.06
3
50
3
4.9
2.4
775
12
20
5
7.7
0.16
0.32
8.5
0.47
4.0
70.36
0.25
16.54
2.32
0.02
0.74
0.43
5.07
3.44
0.10
1.00
100.3
1.4
0.09
63
3.9
1.6
0.9
44
836
297
2
0.51
0.75
6.7
14
1.7
7.4
1.3
0.38
1.1
0.13
0.70
0.11
0.26
0.05
0.41
0.05
4
40
5
4.5
2.7
127
22
18
3
0.54
0.15
1.7
0.28
0.62
6.9
0.42
10
1025
22
35
938
568
6.1
210
<2
37
7
370
28
18
19
0.52
4.8
5
458
67.16
0.43
16.75
3.31
0.03
1.25
2.28
3.63
2.85
0.13
1.42
99.2
1.7
0.08
65
4.3
1.9
2.6
36
1346
291
3
0.38
0.57
6.7
14
1.8
8.6
1.7
0.47
1.4
0.11
0.82
0.14
0.38
0.06
0.33
0.04
6
80
16
6.7
6.1
22
41
20
5
0.35
<0.09
1.9
0.67
88
67.53
0.39
16.53
3.77
0.05
1.52
4.72
3.89
1.05
0.12
0.41
100.0
2.7
0.22
61
6.5
1.7
1.5
19
222
402
3
0.40
1.1
7.5
15
1.8
7.8
1.7
0.56
1.4
0.19
1.1
0.23
0.62
0.09
0.59
0.09
8
68
12
tonalite
6.5
62
38
17
<3
1.3
<0.10
67.54
0.38
16.48
4.01
0.05
1.52
4.67
3.83
1.17
0.11
0.29
100.0
2.0
0.13
51
6.9
1.6
1.1
18
290
396
3
0.58
1.3
7.5
15
1.7
8.0
1.9
0.60
1.4
0.26
1.1
0.20
0.69
0.09
0.65
0.10
8
65
15
5.6
5.2
20
39
17
3
0.45
<0.12
<0.12
0.35
<0.08
0.48
222
82
161
Table A1 (continued)
Sample
E06140 E06144A E06136
Balsapamba batholith
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
162 tonalite
tonalite
tonalite
66.26
0.39
15.80
3.97
0.06
1.62
4.59
3.84
1.74
0.10
0.15
98.5
65.68
0.41
16.31
4.49
0.07
1.65
5.00
3.81
1.16
0.11
0.41
99.1
8.6
7.5
355
4
379
4
66
13
75
15
5
40
17
<3
12
48
18
<3
67.65
0.46
16.99
2.37
0.05
2.02
4.68
3.78
1.29
0.12
0.72
100.1
2.1
0.17
65
3.5
1.7
1.3
22
282
393
3
0.29
0.99
3.5
6.6
0.8
3.5
0.84
0.44
0.81
0.11
0.67
0.12
0.33
0.06
0.36
0.07
11
79
19
3.0
5.7
9
38
16
<3
<0.30
<0.10
E06139
E06138
E06131A E06132 E06135 E06141
Balsapamba porphyry intrusions
E06134
tonalite
qtz-diorite
qtz-diorite
porphyry
67.85
0.46
16.77
2.91
0.05
1.75
4.60
3.80
1.21
0.05
0.41
99.9
1.4
395
2
73
20
8
45
17
<3
0.13
0.47
123
128
14
257
qtz-diorite
porphyry
qtz-diorite
porphyry
qtz-diorite
porphyry
qtz-diorite
porphyry
65.85
0.43
16.10
5.61
0.05
2.00
4.64
3.31
1.27
0.10
0.70
100.0
2.1
0.19
55
9.0
1.6
1.9
27
202
335
2
0.89
1.3
5.5
10
1.2
5.0
0.97
0.40
1.1
0.16
1.2
0.28
0.89
0.14
0.98
0.18
12
97
14
5.8
8.2
27
45
17
<3
0.40
<0.10
57.63
0.64
17.84
7.23
0.06
3.15
5.14
3.12
1.98
0.13
2.76
99.7
1.2
0.06
45
11
1.3
2.2
40
213
344
3
0.46
0.52
4.9
11
1.4
7.4
2.1
0.60
2.2
0.36
1.8
0.34
0.89
0.13
0.77
0.11
14
189
21
24
13
658
49
19
3
1.4
<0.10
62.02
0.65
18.46
3.96
0.07
3.24
4.46
4.02
2.09
0.07
0.84
99.9
2.1
0.14
48
2.4
1.4
2.2
39
287
348
3
0.29
0.33
2.0
4.0
0.5
2.2
0.44
0.39
0.34
0.07
0.38
0.08
0.26
0.04
0.32
0.05
19
138
21
6.6
9.0
69
54
17
<3
1.4
<0.10
66.06
0.43
16.05
5.18
0.08
2.04
4.94
3.42
1.23
0.10
0.44
100.0
2.1
0.17
57
11
1.7
1.8
24
262
333
<2
0.67
1.3
6.3
13
1.8
7.4
1.8
0.51
1.6
0.27
1.8
0.33
1.1
0.15
0.96
0.18
13
98
14
65.62
0.42
16.46
4.48
0.06
1.98
4.67
3.52
1.46
0.11
0.49
99.3
2.3
0.22
56
8.0
1.7
2.6
28
243
348
3
0.92
1.5
8.3
17
2.1
8.2
1.6
0.53
1.0
0.18
1.2
0.24
0.78
0.11
0.81
0.14
12
109
14
11
40
55
16
4
0.97
<0.08
9.9
23
58
17
<3
0.85
0.20
65.63
0.43
16.45
5.70
0.06
2.01
4.68
3.39
1.32
0.10
0.60
100.4
2.2
0.18
59
8.7
1.7
2.5
31
202
340
3
0.98
1.6
5.3
10
1.3
5.0
1.1
0.50
1.1
0.20
1.2
0.30
0.85
0.15
0.98
0.15
13
100
16
5.9
9.6
24
60
16
<3
0.39
<0.11
<0.08
0.84
0.11
0.38
0.19
0.78
0.22
0.78
0.13
0.67
0.18
0.56
1162
61
105
66
79
11392
Table A1 (continued)
Sample
E07044 E07045
Telimbela batholith
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
tonalite
tonalite
E06150
E06149
E06153
Telimbela
porphyry
intrusion
qtz microdiorite
qtz microdiorite
qtz-diorite
porphyry
E05070 E05072 E05073
Gaby porphyry intrusions
E05076
E05078
hbl-plag
porphyry
hbl porphyry
hbl porphyry
64.62
0.55
16.16
5.35
0.10
2.28
4.74
3.36
1.94
0.11
0.90
100.1
4.2
0.29
113
18
4.5
0.9
53
593
309
13
1.8
4.4
16
34
4.3
19
4.8
1.2
4.8
0.75
4.6
0.88
2.8
0.39
2.7
0.39
23
104
24
19
20
31
62
17
4
3.4
0.22
2.4
0.68
65.23
0.50
16.37
5.14
0.07
2.24
5.25
3.59
0.52
0.11
1.04
100.1
2.5
0.15
107
17
2.8
0.7
9
199
349
4
0.90
2.5
10
21
2.6
13
2.6
0.69
2.8
0.45
2.8
0.54
1.7
0.26
1.8
0.29
13
88
25
9.9
12
8
101
17
5
0.27
0.09
1.8
0.80
57.34
0.72
17.58
7.96
0.14
3.86
7.26
3.60
0.82
0.13
0.69
100.1
2.2
0.13
55
13
1.5
2.5
13
202
400
3
0.27
0.52
4.9
12
1.4
7.5
2.0
0.59
2.3
0.38
2.2
0.45
1.3
0.18
1.2
0.18
20
187
36
21
22
69
74
19
5
0.98
<0.15
52.96
0.81
18.03
9.10
0.12
4.70
3.87
3.18
2.79
0.15
3.36
99.1
1.8
0.09
41
15
1.3
18
78
212
321
<2
0.24
0.20
4.7
11
1.4
7.7
2.0
0.65
2.4
0.42
2.5
0.54
1.6
0.20
1.5
0.20
23
248
44
26
23
1857
84
21
105
80
0.42
64.48
0.49
16.47
4.68
0.08
2.02
5.42
3.88
1.02
0.12
0.72
99.4
1.8
0.17
68
7.4
1.9
3.1
18
147
393
3
0.56
0.88
4.3
10
1.4
6.0
1.2
0.51
1.4
0.19
1.3
0.25
0.66
0.11
0.67
0.10
11
114
28
9.7
16
25
44
18
4
70
0.18
63.85
0.52
16.56
3.67
0.06
2.79
7.34
3.94
0.18
0.11
0.66
99.7
2.2
0.18
65
14
1.8
0.4
3
136
309
2
0.42
1.1
4.5
11
1.4
6.5
1.7
0.64
1.8
0.29
2.2
0.46
1.40
0.21
1.53
0.22
21
118
29
9.3
8.9
98
32
16
<3
3.7
<0.17
0.15
0.51
0.33
0.62
0.22
4.2
0.67
1.5
851
143
114
5627
729
1358
hbl-plag
porphyry
62.09
0.52
16.66
6.16
0.06
2.85
6.71
3.78
0.55
0.11
0.30
99.8
14
302
2
136
33
89
29
17
<3
plag-hbl
porphyry
62.63
0.51
16.26
5.27
0.06
2.83
7.17
3.84
0.23
0.11
0.61
99.5
2.1
0.17
65
15
1.8
0.4
4
164
300
2
0.44
1.1
4.1
10
1.4
7.1
1.8
0.65
2.0
0.35
2.3
0.50
1.5
0.23
1.6
0.25
22
136
29
10
8.2
154
32
16
<3
5.1
0.29
64.00
0.54
17.34
2.70
0.05
2.67
6.70
5.20
0.19
0.12
1.04
100.6
14
256
<2
74
30
36
26
15
5
0.78
1.4
216
3785
62.90
0.53
16.24
5.27
0.07
2.93
7.26
3.87
0.17
0.11
0.48
99.8
2.1
0.18
66
15
1.8
0.3
2
134
297
2
0.46
1.1
5.1
12
1.6
6.7
1.9
0.58
1.8
0.32
2.3
0.50
1.5
0.21
1.6
0.24
21
171
30
10
9.6
187
35
16
<3
14
0.27
0.61
1.1
561
713
163
Table A1 (continued)
Sample
E05083a E05083b E05086
Gaby porphyry intrusions
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
164 plag-hbl
porphyry
plag-hbl
porphyry
plag-hbl
porphyry
61.19
0.55
17.05
6.64
0.14
2.38
6.62
3.27
0.60
0.13
0.90
99.5
2.8
0.19
73
17
1.9
0.6
9
425
307
4
0.44
1.0
6.8
14
1.9
8.6
2.1
0.76
2.4
0.38
2.7
0.52
1.7
0.25
1.8
0.27
20
116
24
E05088
E06033
E06041
E06046
E06048
E06050
E06051
E06053
plag-hbl
porphyry
plag-hbl
porphyry
plag-hbl
porphyry
hbl-plag
porphyry
hbl-plag
porphyry
plag-hbl
porphyry
plag
porphyry
+ enclave
plag
porphyry
60.49
0.54
16.52
7.51
0.07
2.57
6.71
4.10
0.30
0.11
0.66
99.6
2.4
0.18
65
12
1.8
0.3
5
96
279
3
0.42
0.98
4.7
11
1.5
6.9
1.8
0.62
2.3
0.31
2.0
0.41
1.2
0.16
1.1
0.20
18
129
31
7.8
7.9
72
32
17
<3
3.8
0.39
61.08
0.53
16.69
7.09
0.07
2.61
6.76
4.19
0.33
0.10
0.66
100.1
2.4
0.16
65
12
1.8
0.3
6
93
283
5
0.44
0.91
5.4
12
1.6
7.8
1.7
0.65
2.3
0.30
2.0
0.39
1.2
0.17
1.2
0.17
18
129
29
8.3
8.7
128
34
17
7
13
0.21
6.1
24
82
17
<3
0.97
<0.12
61.22
0.53
17.21
6.55
0.15
2.34
6.84
3.31
0.67
0.13
1.04
100.0
2.8
0.19
71
16
1.8
0.6
11
413
308
4
0.42
1.0
6.6
14
2.1
8.0
2.0
0.66
2.3
0.37
2.6
0.55
1.7
0.23
1.5
0.29
20
112
24
13
4.6
21
89
17
6
0.56
<0.12
1.2
2.1
0.85
1.8
1.3
0.46
0.95
0.39
2632
2612
169
49
60.80
0.54
17.18
6.52
0.14
2.35
6.77
3.16
0.61
0.11
1.28
99.5
61.60
0.55
17.16
5.43
0.08
2.30
6.70
3.98
0.22
0.12
1.05
99.2
61.07
0.53
16.71
6.53
0.06
2.77
6.58
3.60
0.71
0.10
0.39
99.1
16
17
14
310
5
292
<2
290
<2
116
24
115
17
139
29
19
83
17
4
479
48
17
5
155
30
18
4
62.72
0.52
16.56
5.21
0.06
2.69
7.19
3.92
0.25
0.10
0.52
99.8
2.0
0.12
67
15
1.8
0.4
2
125
293
<2
0.63
1.2
3.6
10
1.4
8.4
2.5
0.74
2.6
0.40
2.3
0.52
1.4
0.23
1.6
0.23
17
133
28
8.3
9.8
440
40
17
5
1.3
0.33
62.80
0.55
16.54
3.33
0.05
2.96
8.08
3.81
0.35
0.10
0.51
99.1
15
304
<2
132
40
79
28
17
<3
0.69
0.74
160
3003
1107
1918
292
61.59
0.53
16.84
5.45
0.04
2.76
7.31
3.45
0.60
0.10
0.45
99.1
2.3
0.17
69
14
1.8
1.1
14
178
294
<2
0.94
1.2
12
20
2.3
8.5
1.7
0.53
1.6
0.28
2.0
0.44
1.4
0.20
1.5
0.25
20
142
23
14
13
822
29
18
<3
17
0.38
60.67
0.57
16.47
6.71
0.06
3.00
7.09
3.39
0.63
0.11
0.43
99.1
2.3
0.16
66
16
1.8
1.4
16
147
287
<2
0.77
1.1
6.2
12
1.6
6.6
1.8
0.58
1.7
0.32
2.1
0.45
1.4
0.22
1.6
0.25
21
156
62
10
11
113
30
18
<3
5.1
0.16
0.39
1.0
0.50
1.6
1765
1906
Table A1 (continued)
Sample
E05090 E06044
Gaby/Papa Grande
porphyry intrusions
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
hbl-plag
porphyry
61.63
0.52
16.87
6.43
0.17
2.84
7.97
2.18
0.18
0.11
1.25
100.2
2.5
0.14
70
13
1.5
0.8
3
75
309
4
0.48
1.1
6.3
14
1.5
7.4
1.6
0.53
1.7
0.29
1.9
0.33
1.2
0.15
1.6
0.19
16
129
35
16
10
35
73
18
5
0.30
<0.17
plag(-hbl)
porphyry
E06042 E06043 E06049
Gaby/Papa Grande tonalite
E06052
E06054
E06055
E06056
tonalite
tonalite
tonalite
tonalite
tonalite
tonalite
62.08
0.50
16.27
6.36
0.12
2.81
7.58
1.95
0.38
0.09
1.26
99.4
61.26
0.54
15.96
6.30
0.07
3.09
6.17
3.46
0.71
0.10
1.47
99.1
63.00
0.51
16.11
5.77
0.07
2.65
6.13
3.80
0.69
0.10
0.43
99.3
12
13
13
299
3
255
<2
283
<2
128
36
147
35
121
29
31
49
17
7
336
39
17
<3
101
46
16
4
6.2
1.5
611
289
8228
94
tonalite
65.27
0.43
15.74
5.21
0.08
2.16
5.26
3.55
1.02
0.09
0.69
99.5
2.5
0.21
69
12
1.9
1.3
22
453
254
<2
1.2
2.8
7.2
14
1.4
6.9
1.6
0.48
1.2
0.29
1.7
0.35
1.1
0.16
1.2
0.18
13
101
23
11
6.5
54
43
16
<3
0.91
0.16
62.51
0.52
16.22
5.77
0.06
2.62
6.73
3.77
0.52
0.09
0.47
99.3
2.7
0.18
75
16
2.1
0.8
10
245
282
2
0.52
1.6
6.3
13
1.7
6.8
1.9
0.56
2.3
0.34
2.4
0.52
1.6
0.22
1.8
0.27
20
125
33
11
9.4
166
34
16
<3
5.32
0.21
0.48
2.3
0.53
1.3
1977
1769
61.84
0.56
16.33
5.07
0.05
2.83
7.34
3.55
0.35
0.11
1.02
99.1
61.90
0.56
16.49
5.25
0.05
2.65
7.15
3.44
0.51
0.10
0.93
99.1
60.21
0.62
16.33
6.68
0.07
2.93
7.25
3.95
0.28
0.11
0.77
99.2
15
15
17
293
<2
299
<2
294
<2
140
24
148
33
162
22
2302
43
17
<3
1095
42
17
<3
775
37
18
4
6691
8838
3256
165
Table A1 (continued)
Sample
E07002 E07005 E06179
Chaucha porphyry intrusions
(Gur-Gur/Tunas)
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
166 dacitic
porphyry
dacitic
porphyry
64.49
0.59
17.57
5.20
0.04
1.51
4.89
3.94
1.20
0.19
0.33
100.0
3.2
0.21
122
8.3
3.1
2.1
40
407
611
4
0.58
1.3
12
25
3.1
13
2.9
0.92
2.6
0.33
1.5
0.31
0.69
0.12
0.74
0.12
9
79
5
4.3
2.0
19
43
22
4
0.61
0.16
1.1
0.12
68.40
0.42
16.55
2.99
0.04
1.23
3.34
4.16
1.75
0.13
1.29
100.3
3.4
0.22
92
7.2
2.5
1.8
37
198
492
7
1.6
5.0
8.0
18
2.2
9.4
2.1
0.64
2.0
0.35
1.9
0.40
1.1
0.19
1.3
0.21
12
62
11
8.7
7.6
34
38
20
5
5.9
0.12
1.7
0.14
478
123
E06181
E07001 E07003
Chaucha batholith
E07008
E06158
E05092 E05093
Quimsacocha
dacite domes
granodiorite
porphyry
granodiorite
porphyry
tonalite
tonalite
qtz-diorite
dacite
73.26
0.19
14.86
1.05
0.04
0.77
3.38
4.35
0.44
0.07
1.34
99.8
2.1
0.31
59
4.6
2.2
1.5
10
189
329
5
1.8
11
5.3
9.4
0.9
4.3
0.94
0.47
0.63
0.11
0.78
0.10
0.45
0.06
0.54
0.11
5
41
7
1.5
4.4
542
40
15
4
2.2
0.23
73.41
0.21
15.10
1.13
0.03
0.92
3.30
4.11
0.41
0.02
1.47
100.1
2.0
0.30
64
1.7
2.4
2.0
10
170
307
4
1.6
10.9
3.6
5.5
0.5
1.7
0.28
0.41
0.33
0.03
0.20
0.06
0.22
0.02
0.25
0.04
5
23
7
1.9
5.2
190
35
15
<3
2.7
<0.23
<0.30
3.6
0.66
3.3
658
371
tonalite
dacite
64.74
0.46
16.57
5.18
0.05
1.96
5.04
2.78
1.23
0.11
1.80
99.9
3.5
0.27
81
13
2.9
2.2
44
359
409
9
2.7
6.5
13
27
3.1
13
2.8
0.75
2.7
0.41
2.2
0.46
1.3
0.18
1.4
0.21
12
89
18
8.2
8.5
9
55
17
4
0.84
0.11
1.2
<0.12
68.65
0.49
16.65
3.10
0.04
2.09
4.04
3.20
1.11
0.10
0.82
100.3
2.7
0.19
77
12
2.3
1.1
42
686
336
5
1.1
2.4
12
25
3.1
13
2.3
0.54
1.9
0.24
1.3
0.23
0.51
0.10
0.64
0.09
7
90
16
6.4
4.0
386
48
17
6
0.34
<0.08
0.95
0.67
65.39
0.50
17.41
4.77
0.04
2.20
4.22
3.17
1.38
0.12
0.85
100.1
3.3
0.26
86
13
2.3
2.2
52
251
356
5
2.2
7.8
10
21
2.5
9.8
2.1
0.63
2.5
0.33
2.2
0.40
1.2
0.16
1.3
0.20
12
111
18
8.0
9.8
818
56
19
<3
2.3
0.41
1.7
0.12
67.87
0.37
16.45
4.30
0.04
1.53
3.79
3.43
0.51
0.09
1.64
100.0
2.3
0.22
63
7.8
1.8
2.9
16
291
416
5
0.90
2.6
5.6
11
1.3
5.1
1.2
0.44
1.1
0.17
1.1
0.21
0.61
0.10
0.66
0.11
8
75
13
2.4
8.6
826
87
17
3
1.1
0.32
69.63
0.35
16.58
2.09
0.02
0.25
2.74
4.26
1.95
0.12
1.20
99.2
2.9
0.21
107
4.8
2.9
0.5
34
876
466
13
0.82
2.4
16
29
3.6
14
2.5
0.63
1.3
0.15
0.79
0.14
0.43
0.05
0.37
0.04
5
36
4
2.3
1.9
14
136
20
<3
0.68
<0.08
2.1
8.9
0.31
0.21
11838
273
1460
954
<3
69.19
0.34
16.37
2.17
0.02
0.64
3.50
4.64
1.80
0.13
0.95
99.8
5.1
858
10
33
3
20
59
20
<3
<3
Table A1 (continued)
Sample
E05094 E05102 E05115
Quimsacocha dacite domes
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
dacite
dacite
69.49
0.34
16.43
2.16
0.03
0.56
3.39
4.67
1.81
0.13
0.75
99.8
4.1
605
11
35
2
20
56
20
5
dacite
69.32
0.37
16.40
2.30
0.02
0.23
2.84
4.39
1.94
0.12
1.10
99.0
3.0
0.19
111
5.7
2.8
0.7
38
888
489
12
0.89
3.3
28
49
5.8
21
3.1
0.70
1.7
0.18
0.91
0.17
0.42
0.07
0.37
0.07
5
35
4
2.3
2.5
10
85
21
3
0.84
0.15
70.33
0.36
16.49
2.14
0.02
0.22
2.66
4.28
2.04
0.12
1.17
99.8
5.1
457
13
39
4
12
84
21
5
0.30
0.20
<3
<3
E06017
E06018
E06019
E06020
E05098
E05099
E05114
Quimsacocha andesite flows
(+subvolcanics?)
E06022
dacite
dacite
dacite
dacite
microdiorite
andesite
68.97
0.35
17.11
2.38
0.02
0.18
2.53
3.96
1.99
0.07
1.71
99.3
2.8
0.17
115
4.6
3.2
1.3
42
984
448
15
0.97
2.7
17
33
3.7
15
2.3
0.54
1.5
0.16
0.81
0.16
0.46
0.04
0.38
0.05
5
32
3
3.1
1.5
19
62
22
<3
0.44
0.20
69.76
0.36
16.59
2.42
0.03
0.15
2.65
4.08
2.03
0.10
1.21
99.4
5.0
463
15
35
4
18
64
21
6
0.25
0.87
<3
<3
<3
68.67
0.38
16.92
2.44
0.02
0.18
3.05
4.55
1.82
0.12
0.97
99.1
3.1
0.18
108
5.0
2.7
0.6
32
794
556
13
0.83
2.2
15
28
3.6
14
2.4
0.63
1.6
0.17
0.81
0.17
0.36
0.06
0.28
0.06
5
38
4
3.0
2.7
21
63
22
3
0.29
0.17
67.42
0.34
16.28
2.17
0.04
0.64
3.37
4.84
1.57
0.12
2.46
99.3
2.8
0.17
103
4.7
2.9
1.4
26
692
649
12
0.69
1.9
15
30
3.8
15
2.4
0.73
1.6
0.17
0.83
0.16
0.38
0.04
0.39
0.05
5
37
4
3.2
3.0
17
79
21
5
1.0
0.18
0.32
0.54
0.52
0.25
<3
9
microdiorite
andesite
62.44
0.68
17.35
4.74
0.09
1.87
5.56
4.06
1.46
0.17
1.19
99.6
2.4
0.14
98
7.8
2.6
0.5
31
666
594
10
0.89
2.4
14
27
3.3
14
2.6
0.81
2.1
0.25
1.2
0.24
0.71
0.10
0.67
0.10
11
127
53
11
5.2
23
81
21
6
1.2
0.24
1.1
0.36
61.60
0.70
17.25
4.76
0.08
2.06
5.26
3.68
1.29
0.17
1.97
98.8
2.5
0.14
90
7.4
2.5
0.4
25
693
578
9
0.63
2.0
14
27
2.9
13
2.6
0.63
2.2
0.24
1.1
0.23
0.59
0.07
0.49
0.06
11
128
12
11
4.3
31
126
21
<3
0.43
<0.15
62.80
0.68
17.43
4.63
0.08
2.23
5.21
4.02
1.41
0.17
1.13
99.8
2.4
0.11
88
7.4
2.4
0.4
30
683
586
9
0.67
2.0
14
28
2.9
13
2.4
0.85
2.3
0.24
1.3
0.25
0.69
0.09
0.44
0.09
10
125
128
10
5.1
19
93
20
<3
0.55
<0.18
61.59
0.77
17.81
5.15
0.05
1.98
5.82
4.53
1.15
0.17
0.42
99.4
2.3
0.13
84
11
2.2
0.6
20
481
696
6
0.52
1.5
13
23
3.2
14
3.0
0.82
2.4
0.32
1.8
0.29
0.90
0.14
0.78
0.14
12
146
19
11
7.7
39
95
23
3
1.4
0.18
<0.24
0.18
0.39
0.25
0.46
0.19
9
32
10
<3
167
Table A1 (continued)
Sample
E06157 E06166 E06172
Saraguro Group at Chaucha
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
168 andesite
andesite
andesite
57.16
0.67
18.49
7.88
0.18
3.00
7.88
2.62
0.75
0.14
0.60
99.4
2.8
0.22
72
20
2.0
1.7
20
220
395
8
0.72
1.8
7.2
15
2.0
9.4
2.5
0.77
2.8
0.43
3.0
0.66
1.9
0.27
1.88
0.33
23
156
10
6.4
9.5
49
93
19
3
0.67
0.05
65.79
0.57
15.62
5.15
0.07
2.26
4.70
2.30
1.64
0.13
1.14
99.4
5.7
0.48
131
20
3.5
5.4
67
534
320
2
2.7
8.1
17
33
4.0
15
3.4
0.80
3.32
0.47
3.1
0.63
2.0
0.27
1.8
0.31
16
122
25
12
17
1082
101
17
5
14
0.59
61.36
0.64
16.74
6.76
0.04
3.20
4.97
2.37
1.59
0.12
1.56
99.4
4.7
0.35
117
19
3.0
9.4
70
266
270
<2
1.3
4.2
10
21
2.6
11
2.8
0.67
3.12
0.40
3.1
0.64
1.8
0.28
1.9
0.29
22
138
30
12
19
508
48
19
4
2.7
0.24
0.23
0.36
0.25
0.89
0.84
2.1
454
1664
308
E06004
E06009 E06010
Saraguro Group at Canicapa
E06011
E06012
E06015
dacite (subvolcanic?)
dacite
(subvolcanic?)
dacite
(subvolcanic?)
andesite
flow
dacite
(flow)
dacite
(flow)
69.86
0.26
14.75
2.36
0.07
0.86
2.83
3.65
2.41
0.10
2.03
99.2
70.59
0.27
15.23
2.47
0.08
0.67
2.58
3.96
2.33
0.12
0.98
99.3
15
13
289
14
340
15
26
2
22
5
8
47
15
4
6
49
16
5
3
10
70.07
0.30
15.28
2.73
0.07
0.81
2.53
3.96
2.37
0.13
0.92
99.2
6.0
0.53
94
12
2.4
1.8
45
723
329
15
1.7
5.3
17
30
3.3
13
2.4
0.68
2.2
0.28
1.7
0.35
0.96
0.15
1.1
0.18
5
27
9
3.5
3.0
11
57
16
4
0.90
0.15
69.37
0.30
15.17
2.76
0.13
1.06
3.14
3.62
2.20
0.13
1.30
99.2
5.9
0.46
100
12
2.5
2.8
42
682
417
15
1.9
5.4
16
30
3.4
13
2.1
0.61
1.9
0.26
1.8
0.36
1.0
0.16
1.2
0.17
5
25
<2
3.7
2.2
7
55
15
5
0.81
<0.10
68.69
0.30
15.60
2.85
0.06
1.02
2.94
3.81
2.28
0.13
1.45
99.1
6.0
0.47
100
12
2.4
2.6
45
702
401
16
1.8
5.5
17
31
3.5
12
2.5
0.59
1.6
0.30
1.9
0.35
1.0
0.16
1.2
0.20
6
26
<2
3.6
<1.5
7
42
17
5
0.86
<0.15
62.03
0.53
16.52
5.27
0.11
2.15
5.22
3.38
1.70
0.14
2.25
99.3
5.1
0.38
87
15
2.3
0.5
36
526
424
5
1.7
4.9
15
28
3.3
13
2.5
0.72
2.5
0.35
2.3
0.46
1.3
0.19
1.2
0.21
11
95
10
11
5.6
13
54
18
4
0.75
0.12
0.38
0.89
0.28
0.56
0.59
0.82
0.12
0.34
<3
11
<3
70
Table A1 (continued)
Sample
E06074 E06075 E06081
Saraguro Group at Portovelo
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
andesite
andesite
57.82
0.72
16.27
7.25
0.16
3.58
7.05
2.52
1.54
0.11
2.08
99.1
59.30
0.70
16.34
6.82
0.13
3.52
6.60
2.62
1.62
0.11
1.69
99.5
19
19
284
8
284
18
159
44
145
40
33
71
19
7
31
68
19
7
286
134
andesite
E06082
E06083A
E06083B
E06084
E06086
E06117
E06118
E06120
andesite
andesite
andesite
andesite
andesite
andesite
andesite
andesite
58.85
0.70
16.16
6.70
0.17
3.52
6.67
2.50
1.48
0.11
2.32
99.2
4.1
0.25
122
20
3.3
1.4
29
677
344
17
1.8
5.6
13
28
3.3
14
3.4
0.72
3.2
0.53
3.4
0.64
2.1
0.30
2.1
0.29
21
146
46
19
19
32
124
18
8
2.0
0.48
57.35
0.70
16.24
7.05
0.21
3.92
6.49
2.67
1.21
0.11
3.50
99.5
3.8
0.24
107
19
3.0
2.1
31
456
299
4
1.4
4.8
12
25
3.0
13
3.1
0.78
3.2
0.49
3.4
0.63
2.1
0.25
1.9
0.30
23
152
47
17
13
26
106
18
6
1.1
0.24
3.8
1.2
1.9
0.76
521
1806
58.48
0.71
16.07
6.40
0.14
3.96
6.06
2.51
1.24
0.12
3.40
99.1
59.37
0.69
15.81
6.38
0.14
2.97
7.26
2.66
0.88
0.12
2.74
99.0
59.11
0.69
16.03
6.44
0.26
3.35
5.59
2.61
2.08
0.12
3.01
99.3
58.40
0.71
15.97
6.49
0.14
3.74
5.24
3.27
1.71
0.11
3.35
99.1
56.13
0.94
19.17
6.94
0.19
3.89
6.68
2.50
0.44
0.17
2.14
99.2
55.09
0.83
19.49
8.13
0.16
4.77
6.00
2.78
0.96
0.15
1.64
100.0
20
20
20
19
21
15
274
8
310
15
280
4
347
7
309
21
350
43
146
46
142
41
139
39
140
54
123
44
142
44
51
128
17
15
11
94
18
22
26
104
18
13
30
92
17
20
39
152
21
14
25
140
23
<3
57.96
0.73
16.39
6.96
0.16
3.47
6.19
2.36
2.10
0.12
2.84
99.3
4.0
0.25
113
20
2.9
5.2
62
552
256
10
1.5
4.8
13
26
2.9
13
2.8
0.76
3.1
0.48
2.9
0.63
1.7
0.25
1.7
0.28
22
163
87
17
17
27
79
18
4
1.7
0.17
1.7
0.98
1722
2498
178
2050
5618
1093
65
169
Table A1 (continued)
Sample
E07013
Saraguro
Group at 3
Chorreras
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
170 granodioritic
porphyry
E06068 E06071 E05-M4
Cangrejos host intrusive
complex
diorite
diorite
60.41
0.47
16.37
5.85
0.06
1.91
3.83
2.14
3.61
0.13
4.47
99.3
3.0
0.21
86
10
2.0
19
109
584
210
3
1.3
2.4
9
19
2.2
9.3
2.1
0.67
1.8
0.31
1.6
0.35
1.1
0.18
1.1
0.16
10
75
9
4.2
4.7
143
51
18
10
3.0
0.33
1.3
19
52.22
0.66
16.07
7.20
0.13
7.94
11.60
1.75
0.48
0.08
1.02
99.2
4.8
0.29
41
18
1.3
0.5
12
67
239
6
0.63
1.8
6
14
2.0
9.1
2.5
0.77
2.6
0.44
3.0
0.61
1.9
0.25
1.6
0.27
45
190
260
30
56
327
101
16
<3
0.95
0.37
55.88
0.38
15.29
6.35
0.10
7.53
9.49
2.15
0.86
0.02
1.04
99.1
5.0
0.35
54
14
1.6
0.6
26
94
229
5
0.71
2.7
7
17
2.1
9.8
2.1
0.72
2.4
0.37
2.4
0.50
1.3
0.22
1.3
0.20
34
154
103
19
32
550
66
16
<3
13
0.80
0.50
9.4
0.37
1.4
168
944
928
qtzdiorite
E06067 E06069 E06070
Cangrejos porphyry intrusions
E06088
plag-hbl
porphyry
plag-hbl
(qtz eyes)
porphyry
qtz-diorite
porphyry
qtz-diorite
porphyry
65.33
0.41
15.35
3.93
0.08
1.96
3.57
2.61
3.36
0.09
3.52
100.2
66.27
0.39
15.04
3.71
0.07
1.76
3.98
2.88
2.60
0.09
3.26
100.1
13
12
201
10
275
11
70
19
65
28
18
59
16
8
13
53
15
4
69
21
E05-M10
plag-hbl
porphyry
plag-hbl
porphyry
61.59
0.60
16.25
6.97
0.04
2.73
5.93
3.27
1.30
0.13
1.20
100.0
4.1
0.24
118
17
2.9
0.9
45
254
302
12
0.87
2.8
9
23
3.0
13
3.2
0.75
3.2
0.43
2.9
0.55
1.7
0.19
1.3
0.24
16
117
37
12
13
388
48
18
6
1.3
0.18
2.3
1.4
64.01
0.55
16.29
4.91
0.03
2.70
5.06
2.86
1.62
0.12
1.47
99.6
4.8
0.29
128
19
3.0
1.8
63
281
285
5
1.7
5.1
19
38
4.3
18
3.9
0.83
3.6
0.49
3.2
0.67
1.8
0.27
1.8
0.25
16
102
42
3.7
12
7
82
17
4
0.34
<0.14
3.2
0.38
61.51
0.63
16.55
5.25
0.03
2.97
5.47
2.89
1.73
0.12
2.03
99.2
4.9
0.36
118
22
3.0
0.9
62
339
317
8
1.2
3.2
19
39
4.8
20
4.4
0.92
4.4
0.58
4.0
0.76
2.1
0.30
1.9
0.28
20
120
40
14
18
483
49
18
6
3.0
0.38
64.89
0.52
15.76
4.11
0.06
2.53
6.81
3.22
0.20
0.10
0.94
99.1
5.1
0.40
115
23
3.0
0.1
3
108
308
7
1.4
4.4
9.3
22
3.0
14
3.2
0.75
3.3
0.53
3.5
0.74
2.2
0.33
2.0
0.32
17
95
34
9.0
12
1028
70
17
7
4.5
0.34
64.40
0.49
16.50
4.51
0.03
2.47
5.08
2.96
1.50
0.10
1.37
99.4
4.7
0.32
113
17
3.0
1.6
50
242
298
4
1.5
4.6
17
35
4.1
16
3.3
0.79
2.6
0.43
2.9
0.57
1.5
0.25
1.7
0.25
15
92
35
2.9
15
8
72
19
<3
0.43
0.64
2.7
0.98
0.42
0.98
0.45
1.9
279
<3
484
2229
16
E06089
Zaruma-Portovelo
porphyry intrusions
Table A1 (continued)
Sample
E06072 E06073 E06090 E06092
Zaruma-Portovelo porphyry intrusions
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
E06112
E06115
E06123 E07023 E06114
Zaruma-Portovelo phaneritic intrusions
E06124
plag-hbl
porphyry
granodiorite
diorite
qtz-diorite
porphyry
qtz-diorite
porphyry
qtz-diorite
porphyry
qtz-diorite
porphyry
64.79
0.40
15.29
3.97
0.08
2.00
4.17
3.13
2.18
0.09
2.92
99.0
3.8
0.33
91
13
2.6
3.4
59
586
313
12
1.9
6.8
13
25
2.7
11
2.1
0.67
2.2
0.29
1.9
0.41
1.2
0.18
1.1
0.21
11
68
20
9.1
8.3
18
55
16
6
0.91
0.37
65.76
0.41
15.86
4.16
0.07
2.12
1.42
3.83
3.70
0.09
1.83
99.3
64.44
0.45
15.97
4.39
0.09
2.19
4.34
3.45
2.06
0.09
2.54
100.0
3.7
0.31
85
12
2.6
2.8
64
520
309
14
2.1
6.5
13
25
2.6
11
2.0
0.59
1.8
0.30
1.9
0.40
1.2
0.16
1.2
0.20
12
80
112
11
12
17
64
16
7
1.0
<0.12
64.67
0.42
15.75
4.23
0.08
2.13
4.53
2.96
2.37
0.09
2.90
100.1
3.8
0.33
92
12
2.5
3.4
75
573
291
12
2.1
6.8
14
27
2.7
11
2.1
0.61
2.0
0.29
2.0
0.37
1.2
0.18
1.2
0.21
12
76
67
11
11
17
56
16
10
0.99
0.19
63.36
0.52
16.55
5.05
0.09
2.57
2.94
4.42
2.47
0.11
2.09
100.2
3.7
0.26
100
14
2.6
4.6
82
518
313
29
1.8
4.8
13
25
2.7
13
2.5
0.67
2.4
0.34
2.3
0.48
1.4
0.21
1.3
0.22
15
98
35
13
15
21
155
17
13
1.5
0.51
65.86
0.51
15.56
4.15
0.11
2.18
4.83
1.83
3.46
0.10
1.18
99.8
4.3
0.32
127
17
3.5
2.7
80
868
228
26
2.4
7.8
14
29
3.4
14
3.1
0.83
3.1
0.40
2.5
0.54
1.6
0.23
1.8
0.28
14
100
28
9.0
8.1
38
83
16
13
1.5
0.50
55.24
0.73
18.79
7.61
0.13
4.37
8.50
2.98
0.55
0.11
0.95
100.0
2.6
0.17
66
14
1.9
1.4
14
239
372
6
0.62
1.7
6.9
16
1.9
8.7
2.4
0.74
2.5
0.39
2.5
0.47
1.4
0.21
1.4
0.20
26
182
34
23
16
68
82
20
<3
0.90
0.23
1.5
1.4
1.5
1.6
4.4
2.2
3.5
2.0
0.56
0.66
14
41
47
562
963
12
293
9
76
25
13
52
16
7
1.3
1.1
11
<3
diorite
diorite
qtz-diorite
porphyry
60.46
0.67
17.41
6.42
0.10
3.00
5.60
2.98
1.56
0.12
1.25
99.6
3.8
0.24
108
19
3.1
4.5
51
463
329
22
1.4
4.9
11
24
3.0
13
3.2
0.72
3.4
0.52
3.2
0.61
1.8
0.28
1.7
0.26
19
121
26
15
13
9
98
18
11
0.93
0.39
1.3
2.4
62.76
0.47
14.52
4.20
0.08
1.86
3.33
2.35
2.60
0.09
7.90
100.2
57.47
0.70
16.11
7.28
0.14
4.60
4.74
2.27
1.84
0.13
3.77
99.1
19
18
308
20
228
14
91
27
158
90
25
65
16
11
25
92
17
16
173
2148
4411
171
Table A1 (continued)
Sample
E07016 E07017
El Mozo intrusions
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
172 dioritic
porphyry
dioritic
porphyry
E07020
E07026 E07029
Curiplaya intrusions
E07027
E07031
E07028
plag-hbl
porphyry
plag-hbl
porphyry
microdiorite
qtz dioritic
porphyry
hbl-plag
porphyry
plag porphyry
(Celica flow?!)
62.40
0.58
16.83
6.99
0.17
3.50
1.93
2.15
0.91
0.14
3.72
99.3
4.6
0.32
85
16
2.3
10.0
28
249
197
8
1.9
5.2
15
30
3.5
14
3.0
0.83
2.6
0.45
2.6
0.47
1.5
0.25
1.5
0.24
15
98
17
6.0
4.5
<2
193
18
<3
1.1
0.15
1.2
<0.12
62.41
0.57
17.53
6.46
0.24
2.79
3.74
2.56
1.22
0.14
2.54
100.2
5.0
0.41
87
19
2.3
7.2
28
552
316
29
2.0
5.4
16
31
3.5
15
3.4
0.93
3.2
0.49
3.0
0.59
1.9
0.24
2.0
0.27
16
105
341
13
5.5
7
233
18
4
1.2
0.33
1.3
0.21
61.31
0.56
17.02
6.01
0.13
2.48
5.67
2.81
1.70
0.16
1.19
99.0
5.2
0.41
91
25
2.5
1.5
45
558
459
10
2.0
5.4
20
32
4.1
17
3.7
1.1
3.9
0.52
3.5
0.71
2.3
0.35
2.0
0.37
15
94
13
11
5.1
15
53
18
5
0.67
0.14
1.5
<0.14
64.73
0.31
14.72
3.21
0.25
0.98
4.96
1.87
2.46
0.11
6.43
100.0
3.0
0.17
89
13
2.4
3.0
60
1233
142
16
1.3
5.7
18
34
3.7
15
2.9
0.75
2.2
0.30
1.8
0.37
1.2
0.17
1.2
0.21
6
43
<2
5.4
1.4
20
160
15
<3
0.64
0.16
1.0
3.9
50.59
0.86
17.30
8.77
0.19
3.95
7.91
3.35
0.38
0.23
5.89
99.4
2.1
0.14
53
18
1.5
1.8
9
277
516
12
0.77
2.8
13
27
3.4
16
3.7
1.1
3.4
0.48
3.0
0.53
1.9
0.23
1.8
0.27
18
195
9
20
7.7
8
77
18
<3
0.31
<0.07
1.0
1.5
62.18
0.51
16.48
4.98
0.14
2.76
2.18
5.43
1.78
0.17
2.44
99.1
2.5
0.12
65
15
1.9
0.8
34
1267
637
4
0.75
2.5
14
30
3.7
16
3.3
0.99
2.8
0.40
2.5
0.44
1.5
0.20
1.5
0.21
11
79
13
8.9
5.2
20
51
17
13
0.26
0.09
1.1
3.0
63.17
0.64
16.10
5.35
0.11
2.11
5.60
3.97
0.44
0.17
1.65
99.3
1.9
0.09
85
18
2.3
0.4
5
387
432
4
0.33
1.3
7
17
2.3
12
2.8
0.94
3.2
0.51
2.9
0.62
1.7
0.22
1.9
0.27
10
64
8
6.9
2.1
147
38
18
5
<0.20
0.14
2.2
0.61
56.18
0.69
17.49
7.25
0.18
2.82
5.81
4.79
0.48
0.23
3.24
99.2
3.0
0.14
76
21
2.1
2.1
9
293
193
9
1.0
3.4
17
36
4.4
19
4.1
1.4
3.9
0.53
3.4
0.66
2.0
0.28
1.9
0.30
16
102
5
17
5.7
55
57
19
10
0.50
0.12
0.99
0.92
503
6359
179
105
390
168
216
373
Table A1 (continued)
Sample
E05122 E05123
Basement units
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
Macuchi subvolcanic at
Junin
64.66
0.64
15.66
3.22
0.02
2.40
0.20
1.74
7.49
0.13
1.19
97.4
25
56
<2
169
51
3864
30
18
<3
1290
E06148
E06145
Macuchi basalt at
Balsapamba
E07040
E07042
E07043
Macuchi subvolcanic at Echeandia
63.95
0.59
16.63
5.44
0.05
2.87
1.21
2.67
3.84
0.21
2.14
99.6
3.0
0.15
106
20
3.2
2.3
66
262
199
<2
2.4
5.0
11
23
2.9
14
3.0
0.62
3.1
0.51
3.2
0.69
2.1
0.27
1.9
0.32
14
150
37
12
18
346
61
18
<3
41
0.21
50.77
0.93
17.45
11.27
0.29
5.42
9.17
3.17
0.98
0.17
0.35
100.0
1.0
0.05
66
25
2.2
0.9
13
120
318
<2
0.42
0.83
5.0
12
2.0
9.6
2.6
0.83
3.9
0.57
4.1
0.90
2.5
0.38
2.5
0.40
49
386
49
33
17
678
124
19
<3
0.67
<0.12
52.81
0.94
17.42
10.71
0.58
4.54
6.90
3.89
1.34
0.24
0.51
99.9
1.0
0.06
66
25
2.0
1.7
27
544
616
9
0.23
0.81
5.3
12
1.9
10
2.7
0.90
3.8
0.59
4.1
0.88
2.5
0.36
2.6
0.43
47
337
51
29
26
43
140
17
3
0.63
0.12
0.17
5.5
0.44
0.35
0.28
0.44
459
564
94
E06144B
xenolith in
Balsapamba
tonalite (Macuchi?)
49.45
0.93
21.23
9.10
0.19
4.97
7.95
2.51
0.12
0.14
2.55
99.1
1.0
0.06
44
16
1.2
0.2
2
134
288
5
0.22
0.63
4.1
10
1.5
7.9
2.6
0.89
2.8
0.39
2.7
0.59
1.8
0.20
1.6
0.26
33
313
38
21
22
64
143
20
5
0.30
<0.12
1.81
0.59
56.04
0.72
17.54
10.61
0.27
5.35
1.33
4.87
0.16
0.09
3.14
100.1
0.6
0.06
34
25
1.0
0.2
2
34
108
4
0.18
0.36
2.9
8
1.2
6.1
2.1
0.75
3.2
0.55
3.7
0.81
2.4
0.29
2.0
0.30
30
288
78
14
5.2
338
135
17
<3
0.32
0.14
1.5
0.34
57.72
0.75
17.25
7.71
0.13
3.89
7.48
2.93
1.45
0.13
-0.04
99.4
1.6
0.13
88
18
2.6
2.5
32
268
374
6
0.87
1.9
6.3
15
2.2
11
3.0
0.77
3.2
0.51
3.0
0.60
1.7
0.28
1.8
0.28
25
175
38
20
22
97
66
19
8
1.7
0.10
1.6
0.79
49.02
0.72
16.37
16.82
0.28
4.58
6.86
3.97
0.97
0.14
0.13
99.9
2437
5711
145
263
34
332
8
309
162
14
113
24
5
173
Table A1 (continued)
Sample
E06213
E06035
E06036
Basement units
Lithology
SiO2
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
Total
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
Pb
U
Th
La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
Sc
V
Cr
Co
Ni
Cu
Zn
Ga
As
Mo
Ag
Sn
Sb
W
S
174 Rio Cala
subvolc. at
Junin
CCOP basalt at Gaby
53.53
1.08
18.05
8.39
0.06
7.52
0.30
0.41
7.00
0.11
2.56
99.1
3.1
0.18
60
18
2.1
4.8
120
162
25
<2
1.6
1.6
7.0
16
2.1
9.8
2.5
0.55
3.0
0.45
3.0
0.62
1.8
0.24
1.7
0.28
35
241
335
28
94
1515
64
21
<3
5.9
0.16
1.7
0.70
50.65
2.01
13.07
14.86
0.21
5.36
9.71
1.60
0.15
0.16
1.39
99.2
5.8
0.35
97
36
2.9
0.2
2
70
105
<2
0.15
0.50
5.8
15
2.0
12
3.7
1.3
5.6
0.85
5.9
1.2
3.8
0.54
3.7
0.57
44
482
84
38
55
159
175
20
5
0.80
0.24
295
1663
49.73
2.11
13.20
15.55
0.17
6.36
8.50
1.62
0.13
0.17
1.71
99.3
40
80
<2
489
69
145
154
21
6
1.6
0.43
1020
Tab. A2: Relative changes in concentration compared to least-altered reference sample concentration calculated
from Grant (1986); isocon constructed assuming constant mass of Al2O3
Sample
Magmatic
center
Lithology
Alteration
SiO2
TiO2
Al2O3
Fe2O3
E06200*
Junin-Cuellaje
Junin-Cuellaje
granodiorite
granodiorite
propy
K + propy
66.22
0.41
16.58
3.60
E06202
-1%
6%
4%
E06206A
Junin-Cuellaje
granodiorite
K
4%
6%
9%
13%
7%
5%
10%
4%
E06205A
Junin-Cuellaje
granodiorite
K
3%
-2%
3%
13%
9%
4%
20%
-4%
E06206B
Junin-Cuellaje
granodiorite
K
11%
1%
1%
3%
-3%
7%
25%
1%
E06135*
Balsapamba
Balsapamba
qtz-diorite por.
qtz-diorite por.
K + propy
K + propy
65.63
0.43
5.70
2.01
4.68
3.39
1.32
0.10
E06131A
3%
3%
-7%
4%
8%
4%
-4%
3%
E06132
Balsapamba
qtz-diorite por.
K + propy
0%
-2%
-21%
-1%
0%
4%
11%
10%
E06141
Balsapamba
qtz-diorite por.
K + propy
3%
2%
1%
2%
1%
0%
-2%
2%
16.45
CaO
Na2O
K2O
P2O5
1.69
3.99
12%
10%
4.42
1.46
0.14
1%
3%
1%
E05078*
Gaby
hbl por.
Na-Ca
62.90
0.53
5.27
2.93
7.26
3.87
0.17
0.11
E05070
Gaby
hbl-plag por.
Na-Ca
0%
-4%
-32%
-6%
-1%
0%
4%
-2%
E05073
Gaby
plag-hbl por.
Na-Ca
-1%
-4%
0%
-3%
-1%
-1%
35%
0%
E05083a
Gaby
plag-hbl por.
Na-Ca
-5%
0%
40%
-14%
-9%
4%
73%
-2%
E05083b
Gaby
plag-hbl por.
Na-Ca
-5%
-3%
31%
-13%
-10%
5%
89%
-12%
E05086
Gaby
plag-hbl por.
Na-Ca
-7%
0%
20%
-23%
-13%
-20%
240%
12%
E05088
Gaby
Gaby
plag-hbl por.
hbl-plag por.
propy
Na-Ca
-8%
-4%
17%
-25%
-11%
-19%
275%
11%
-2%
-4%
-3%
-10%
-3%
-1%
44%
-11%
E06048
16.24
MgO
E06051
Gaby
plag por.
K
-6%
-4%
0%
-9%
-3%
-14%
240%
-12%
E06053
Gaby
plag por.
K
-5%
6%
25%
1%
-4%
-14%
266%
-1%
E05090
Gaby
hbl-plag por.
-6%
-4%
17%
-6%
6%
-46%
1%
-4%
-2%
Na-Ca/propy
E05072
Gaby
hbl-plag por.
Na-Ca/propy
-4%
-4%
14%
-5%
-10%
-5%
216%
E05076
Gaby
Gaby
hbl por.
plag-hbl por.
Na-Ca
propy
-5%
-3%
-52%
-14%
-14%
26%
4%
2%
-9%
-3%
17%
-24%
-12%
-23%
242%
-6%
E06041
Gaby
plag-hbl por.
propy
-7%
-1%
-3%
-26%
-13%
-3%
22%
3%
E06046
Gaby
hbl-plag por.
K + Na-Ca
-6%
-3%
20%
-8%
-12%
-10%
306%
-12%
E06050
Gaby
plag-hbl por.
K + Na-Ca
-2%
2%
-38%
-1%
9%
-3%
102%
-11%
E06044
Gaby
plag(-hbl) por.
Na-Ca/propy
-1%
-5%
20%
-4%
4%
-50%
121%
-19%
E06081*
Portovelo
Portovelo
andesite
andesite
propy
propy
58.85
0.70
6.70
3.52
6.67
2.50
1.48
0.11
E06082
-3%
0%
5%
11%
-3%
6%
-19%
-2%
E06120
Portovelo
andesite
propy (weak)
-3%
2%
2%
-3%
-9%
-7%
39%
7%
E06074
Portovelo
andesite
propy
-2%
2%
8%
1%
5%
0%
3%
0%
E06075
Portovelo
andesite
propy
0%
-2%
1%
-1%
-2%
4%
8%
0%
E06083A
Portovelo
andesite
propy (strong)
0%
2%
-4%
13%
-9%
1%
-16%
9%
E06083B
Portovelo
andesite
propy (strong)
3%
0%
-3%
-14%
11%
9%
-40%
11%
E06084
Portovelo
andesite
propy
1%
-1%
-3%
-4%
-16%
5%
42%
9%
Portovelo
andesite
propy
0%
3%
-2%
8%
-21%
32%
17%
0%
E06033
E06086
16.16
Abbreviation key minerals: qtz - quartz, hbl - hornblende, plag - plagioclase; alteration: propy - propylitic, K - potassic, Na-Ca - sodiccalcic (cf. Seedorff et al. 2005); por – porphyry intrusion.
* reference sample where concentration is given in wt.% (major element oxides) and ppm (trace elements)
175
Tab. A2 (continued)
Nb
Ta
Zr
Y
Hf
Cs
Rb
Ba
Sr
U
Th
Sample
E06200*
2.4
0.2
78.8
5.9
2.1
1.8
29.5
1406
563
1.0
1.7
E06202
5%
-8%
-9%
27%
-3%
-31%
-14%
-66%
3%
-23%
-13%
11%
-1%
7%
24%
9%
-43%
-4%
-52%
2%
-15%
35%
E06206A
E06205A
20%
0%
E06206B
35%
-7%
E06135*
2.2
0.2
59.0
8.7
1.7
2.5
30.7
202
340
1.0
1.6
E06131A
-3%
-6%
-1%
26%
1%
-23%
-21%
33%
0%
-30%
-13%
E06132
2%
18%
-5%
-8%
-2%
4%
-8%
20%
2%
-7%
-3%
E06141
-7%
5%
-5%
6%
-4%
-20%
-12%
2%
1%
-8%
-13%
E05078*
2.1
0.2
66.2
14.6
1.8
0.3
2.0
134
297
0.5
1.1
E05070
0%
-3%
-4%
-7%
3%
22%
30%
-1%
2%
-10%
-5%
E05073
-1%
-5%
-3%
1%
4%
13%
99%
22%
1%
-3%
2%
E05083a
9%
-1%
-3%
-18%
-1%
3%
122%
-30%
-8%
-10%
-14%
E05083b
10%
-15%
-5%
-17%
-1%
7%
196%
-33%
-7%
-7%
-20%
E05086
25%
2%
4%
9%
5%
74%
325%
201%
-2%
-8%
-11%
-13%
E05088
21%
3%
1%
5%
-5%
94%
415%
190%
-2%
-14%
E06048
-10%
-35%
0%
1%
0%
9%
18%
-9%
-3%
36%
3%
E06051
5%
-11%
0%
-6%
-3%
232%
569%
28%
-5%
99%
5%
E06053
4%
-14%
-2%
5%
2%
344%
671%
8%
-5%
66%
2%
E05090
10%
-23%
1%
-15%
-20%
153%
67%
-46%
0%
0%
-4%
E05072
-7%
-1%
E05076
-13%
-19%
E06033
4%
-1%
E06041
10%
-7%
E06046
-7%
-5%
E06050
-2%
0%
E06044
-17%
0%
E06081*
4.1
0.3
121.8
19.7
3.3
1.4
28.5
677
344
1.8
5.6
E06082
-7%
-6%
-12%
-7%
-8%
46%
10%
-33%
-13%
-21%
-16%
E06120
-3%
-1%
-9%
-2%
-14%
259%
115%
-20%
-27%
-18%
-17%
E06074
-2%
-18%
E06075
-4%
-18%
E06083A
1%
-20%
E06083B
4%
-8%
E06084
1%
-18%
E06086
-1%
2%
176 Tab. A2 (continued)
La
Nd
Sm
Eu
Gd
Dy
Yb
Sc
V
Cr
Ni
Sample
E06200*
8.5
9.0
1.9
0.5
1.5
0.9
0.5
7.9
81
27
13.3
E06202
7%
8%
-1%
-2%
-3%
43%
20%
10%
4%
4%
32%
31%
18%
8%
13%
40%
40%
62%
20%
26%
E06206A
3%
53%
E06205A
-12%
-1%
E06206B
-8%
19%
E06135*
5.3
5.0
1.1
0.5
1.1
1.2
1.0
12.5
100
16
9.6
E06131A
23%
52%
64%
4%
47%
49%
1%
8%
0%
-10%
16%
E06132
58%
64%
46%
7%
-7%
-6%
-17%
-5%
9%
-13%
3%
E06141
6%
2%
-12%
-19%
4%
1%
2%
2%
-1%
-11%
-13%
E05078*
5.1
6.7
1.9
0.6
1.8
2.3
1.6
21.3
171
30
9.6
E05070
-12%
-4%
-13%
8%
-4%
-9%
-4%
-2%
-32%
-5%
-8%
E05073
-19%
6%
-5%
11%
8%
-1%
5%
1%
-21%
-3%
-14%
E05083a
-9%
1%
-6%
5%
21%
-14%
-30%
-15%
-26%
2%
-18%
E05083b
3%
14%
-15%
8%
22%
-16%
-22%
-18%
-27%
-6%
-12%
E05086
26%
23%
6%
24%
21%
8%
12%
-12%
-35%
-24%
-40%
E05088
22%
13%
-3%
6%
20%
6%
-11%
-13%
-38%
-24%
-54%
E06048
-30%
24%
30%
24%
37%
-2%
-2%
-21%
-24%
-8%
0%
E06051
121%
22%
-16%
-13%
-17%
-20%
-5%
-8%
-20%
-26%
29%
E06053
20%
-2%
-8%
-2%
-11%
-12%
2%
-5%
-10%
104%
18%
E05090
19%
6%
-21%
-12%
-10%
-21%
-3%
-26%
-27%
12%
1%
E05072
-22%
7%
E05076
-59%
-6%
E06033
-36%
-24%
E06041
-36%
-46%
E06046
-21%
-6%
E06050
-24%
31%
E06044
-25%
20%
E06081*
13.3
13.8
3.4
0.7
3.2
3.4
2.1
20.9
146
46
19.4
E06082
-12%
-8%
-9%
7%
-1%
-1%
-10%
8%
4%
2%
-31%
E06120
-7%
-6%
-19%
4%
-6%
-18%
-20%
4%
-13%
10%
87%
E06074
8%
-5%
E06075
-2%
-14%
E06083A
1%
1%
E06083B
-1%
-9%
E06084
-4%
-15%
E06086
-3%
19%
177
178 CHAPTER VI GENERAL CONCLUSIONS AND OUTLOOK Conclusions The main objective of this thesis was to explore the mutual relationships of geodynamic envi‐
ronment, the geochemical features of arc mag‐
matism, and porphyry‐related ore deposit forma‐
tion in Ecuador at a regional scale. Organized in four principal chapters, I describe and discuss new geochronologic data on arc magmatism (Chapter 2) and porphyry‐related hydrothermal systems in Ecuador (Chapter 3), as well as on the isotopic (Chapter 4) and major and trace element geochemical composition (Chapter 5) of Late Ter‐
tiary porphyry‐related arc magmas. The geody‐
namic evolution of the Panama basin and the Ec‐
uadorian margin as known from published data is referenced where appropriate, and supple‐
mented by a recalculation of the most recent set of available Farallon/Nazca‐South America plate convergence parameters for central Ecuador (Chapter 2), a discussion of the oblique subduc‐
tion system of Ecuador and its implications for crustal strain partitioning (Appendix of Chapter 2), and an updated estimate of the collisional tim‐
ing of the Carnegie Ridge seamount chain with the Ecuadorian margin (Chapter 3). As shown in Chapter 2, the regional distribution pattern of Tertiary arc magmatism in Ecuador shows a strong dependency on crustal structures (often reactivated suture zones) and slab dip, the latter probably influenced by the subduction of buoyant oceanic features such as the Inca pla‐
teau (southern Ecuador‐northern Peru) and the Carnegie Ridge seamount chain (northern Ecua‐
dor). An arc magmatic flare‐up event comprising widespread ignimbrite eruption and batholith construction occurred in the Late Oligocene to Early Miocene and coincides in time with an ac‐
celeration of Farallon/Nazca‐South America con‐
vergence rates. Mid‐ to Late Miocene slab flat‐
tening progressively shut down arc magmatism in the southern Ecuadorian (and northern Peruvian) arc segment, and moderate slab shallowing in the Late Miocene led to the eastward migration of arc magmatic activity in northern Ecuador. Chapter 3 demonstrates that porphyry‐related ore deposits in Ecuador formed throughout the Miocene (and latest Oligocene) and follow the general distribution of arc magmatism in space and time. These ore deposits share many charac‐
teristics with, and may be regarded as the north‐
ern extension of the central‐northern Peruvian Miocene metallogenic belt of major economic importance. Ore deposits are often located close to regional structures where focused upper crustal magma ascent is suggested by the occur‐
rence of batholith‐scale intrusive clusters. Por‐
phyry‐related ore deposits tend to form towards the final stages of batholith assembly. Suitable exposure (in particular, the lack of younger volcanic cover sequences in the Mid‐ to Late Miocene flat slab region of southern Ecua‐
dor) and preservation levels are key factors in controlling the outcropping parts of porphyry systems and their total tonnages (compare, for example, the deeply eroded western foothills of the Western Cordillera where locally only the roots of porphyry systems are preserved, and the bulk of the porphyry Cu mineralization has often been eroded). A first‐order spatio‐temporal cor‐
relation between ore deposit formation and seamount chain subduction, as proposed by some authors for the central and southern An‐
des, is not observed in Ecuador. However, by its influence on slab dip (with a potential lag time of several m.y.), the subduction of buoyant oceanic features may strongly influence the outcrop pat‐
tern of arc magmatism, and thus the exposure and preservation levels of older porphyry‐related ore deposits. Pronounced crustal evolution of Tertiary arc magmas can be demonstrated by systematically changing isotopic (Sr, Nd, Pb) compositions across the arc, reflecting variable basement units at depth (Chapter 4). Entirely oceanic basement 179
domains host both Au‐Cu and Cu‐Mo porphyry systems suggesting that crustal basement com‐
position does not control the type of mineraliza‐
tion encountered in a given porphyry system in Ecuador. As discussed in Chapter 5, trace element compo‐
sitions of Ecuadorian arc magmas change sys‐
tematically through time suggesting progressive crustal thickening, downwards migration of the focus of crustal magma evolution (i.e., hot zones), or a combination of both. Moreover, local factors such as different melt water contents may influ‐
ence variations in trace element trends through time. Additional variations in the magma source composition (e.g., with the change from Farallon to Nazca seafloor subduction at the Ecuadorian trench during the Miocene) are possible, but can‐
not be evaluated with the current dataset of mostly highly differentiated compositions. With few exceptions, Late Tertiary Ecuadorian arc magmas are mainly non‐adakitic until the Late Miocene when adakite‐like signatures are ob‐
served frequently on a regional scale, a trend that dominates until the present day. Chapter 5 shows that, for the most part, the geo‐
chemical signatures of porphyry intrusions, in particular adakite‐like features, are similar to the regional geochemical characteristics of arc mag‐
matism for a given period of time. Distinct geo‐
chemical signatures of porphyry intrusions with respect to spatially associated phaneritic intru‐
sions are occasionally observed. In the latter case, however, there is a time gap of several m.y. between mineralizing porphyry intrusion and phaneritic pluton emplacement, and the change in geochemical composition through time occurs on a regional scale. These observations suggest that parental melts of Tertiary porphyry‐related ore deposits in Ecuador are related to ordinary arc magmatism; they are not associated with a distinctive petrogenetic source process such as slab melting requiring a special geodynamic envi‐
ronment. This agrees well with multiple studies which have shown that adakite‐like features of active Ecuadorian arc volcanoes are the product of crustal magma evolution, rather than being exclusively related to a specific petrogenetic source process such as slab melting. As noted above, however, the post‐porphyry establish‐
ment of a flat slab segment may locally create 180 optimum exposure and preservation conditions for Miocene porphyry‐related ore deposits in southern Ecuador (and northern Peru). From a regional mineral exploration point of view, the following major conclusions apply for Ecuador: 
Miocene intrusive rocks in Ecuador are inti‐
mately associated with porphyry‐related ore deposits at structurally favorable sites. 
The cluster of southern Ecuadorian porphyry‐
related ore deposits can be regarded as the northward extension of the Miocene metal‐
logenic belt of northern‐central Peru. The belt is possibly continuous towards Colombia. 
A special geodynamic setting (e.g., ridge sub‐
duction) facilitating slab flattening and a shutdown of arc magmatism may create fa‐
vorable exposure conditions for porphyry‐
related ore deposits formed earlier. This is the case for Miocene ore deposits in south‐
ern Ecuador where Late Miocene slab flatten‐
ing occurred. 
Miocene porphyry‐related ore deposits in Ecuador are not preferentially associated with intrusions of special (in particular, ada‐
kite‐like) geochemical compositions. Por‐
phyry parental melts are not petrogenetically related to slab melting, and thus do not re‐
quire a special geodynamic setting to form, although the latter may strongly influence the exposure conditions of porphyry‐related mineralization as noted above. Outlook This thesis was designed to discuss the metal‐
logenic and petrogenetic evolution of multiple arc segments for which available literature data are very sparse. Consequently, the conclusions of this work have to be rather general, and several problematic issues arise; each of these was partly addressed in the discussion of the appropriate chapter(s) of this thesis, but will briefly be re‐
addressed in the following, combined with a gen‐
eral outlook for future work directions. Three of the main issues are: (1) The regional aspect of the subject of this thesis implies that, in addition to newly acquired data within the frame of this PhD project, literature data had to be taken into account for various dis‐
cussions (in particular with respect to the spatio‐
temporal distribution of arc magmatism). Are geochronologic data used from the literature ro‐
bust at a scale significant for this work? The lack of robust geochronologic data in Ecua‐
dor often necessitated the consideration of pub‐
lished ages based on the potentially disturbed K‐
Ar isotopic system for the purpose of a regional discussion of arc magmatism. Where robust geo‐
chronologic data were available for the same lithology, existing K‐Ar data show broadly (on a multi‐m.y. scale) similar ages. All K‐Ar ages used in this thesis were screened carefully and a sys‐
tematic disturbance was not detected; single po‐
tentially disturbed K‐Ar ages were excluded from the database (Appendix Chapter 2). A systematic age bias and resulting significant inaccuracies in the discussion (especially of Chapter 2) are thus unlikely, although they cannot be entirely ruled out. Future geologic studies in Ecuador aiming at constructing a regionally extensive and robust geochronologic framework, both of Tertiary and older lithologies, would be highly desirable and seem scientifically and economically justified, given the geologic complexity and metallogenic potential of the country. (2) How representative (in terms of space‐time distribution and geochemical composition) is the regional sampling approach of this study for Ec‐
uadorian arc magmatism on a regional scale, and for individual ore deposits on a local scale? This thesis does not comprehensively assess the geochemical evolution of Tertiary arc magma‐
tism. Instead, it focuses on intrusive suites asso‐
ciated with the major porphyry‐related ore de‐
posits known during the early planning stages of this PhD project (2005‐2007). Several major batholiths (such as Portachuela in southern Ecua‐
dor) which are partly associated with Tertiary porphyry‐related mineralization (e.g., the Rio Blanco porphyry at the Peruvian‐Ecuadorian bor‐
der) had to be excluded completely due to diffi‐
cult field logistics; arc volcanics were generally not sampled except on few occasions where they are spatially associated with epithermal minerali‐
zation. Consequently, the trace element vs. time distribution trends shown in Chapter 5 may not necessarily be equivalent to (or representative of) the general evolution of Tertiary arc magma‐
tism for which very limited reliable data exist. Thus, although present‐day Northern Volcanic Zone (NVZ) geochemical data are used for com‐
parative purposes in Chapter 5, the discussed systematic differences between NVZ and earlier Tertiary magmatism might in part be due to in‐
complete coverage of the latter, in particular as porphyry‐related intrusions were often of highly differentiated compositions. Sampling at a given deposit site was not carried out using a systematic grid, but instead was con‐
trolled by outcrop and drill core accessibility, al‐
though an attempt was made to include all rele‐
vant lithologies in the sampling campaign. As knowledge of the individual ore deposit geology at the time of sampling, depending on the devel‐
opment stage of a given deposit, was highly vari‐
able, it is further possible that sampling was somewhat biased and not all lithologies were in‐
cluded (this applies to both geochemistry and geochronology). Hence, the number of samples for geochemical analysis for a given ore deposit does not reflect the volume of a particular lithol‐
ogy. Samples chosen for geochronologic analysis provide snapshots of the temporal evolution of the magmatic‐hydrothermal system, but do not cover its complete history, especially when mul‐
tiple porphyry intrusions are present. This issue complicates comparing ages obtained from dif‐
ferent isotopic systems (e.g., U‐Pb zircon vs. Re‐
Os molybdenite). More extensive studies of some of these ore deposits may only prove useful once outcrop accessibility improves, and significant progress in project development on behalf of the concession holder has been made. (3) What is the significance of the observed geo‐
chemical trends? I attempt to statistically evaluate hydrothermal alteration‐induced element mobility in Chapter 5 beyond the generic "LILE are mobile, HFSE are immobile" scheme by using isocon plots where possible. The results of this exercise are some‐
what discouraging as most elements (including HFSE such as Y, Nb, and Ta) may be variably mo‐
bile, and scatter within a given lithology often exceeds the expected 10% relative error as in‐
ferred from analytical precision. Although some 181
elements, in particular Sr, do not seem to be sig‐
nificantly affected by alteration in three out of four isocon plots shown in Chapter 5, the as‐
sumption that these isocon plots are representa‐
tive for the whole dataset cannot be proven. However, individual trace elements and trace element ratios mostly show correlated behavior where expected (e.g., amphibole fractionation causes increasing Sm/Yb and decreasing Y, both of which is observed in the dataset on a regional scale), and trace element distributions through time are consistent with the conclusions pre‐
sented in Chapter 5 (cf. synthesis above). There‐
fore, while alteration has undoubtedly affected trace element concentrations in most samples to a significant extent, broad regional trends seem to hold some petrogenetic relevance. Further‐
more, as (according to isocon plots) both Sr and Y are only moderately mobile in most cases, I ex‐
pect that a genuine distinction of adakite‐like vs. non‐adakitic features of porphyry and phaneritic intrusions is possible (with the exception of the Junin porphyry system where strong phyllic al‐
teration caused massive feldspar replacement by sericite, and measured Sr contents probably do not approximate fresh rock values). For future works, high‐resolution geochemical studies at the deposit scale might allow quantify‐
ing petrogenetic trends for suites of cogenetic samples, and to better account for alteration ef‐
fects by more detailed mineralogical studies (e.g., electron microprobe analysis of igneous and al‐
teration mineral compositions) which may be used as a reference for regional studies under comparable conditions in a given arc segment. As most investigated porphyry systems comprise multiple intrusive phases, such a high‐resolution study would ideally be coupled with more de‐
tailed geochronologic work based on a solid field campaign which should also include local‐
regional structural studies. The latter is of major importance when evaluating geodynamic changes and their feedback on arc magmatism as it might help unraveling the transitions from in‐
ferred regional to observed local stress field variations, and the potential implications for crustal magma evolution and emplacement. 182