THE GENESIS OF ‘GIANT’ COPPER‐ZINC‐GOLD‐SILVER VOLCANOGENIC MASSIVE SULPHIDE DEPOSITS AT TAMBOGRANDE, PERÚ: AGE, TECTONIC SETTING, PALEOMORPHOLOGY, LITHOGEOCHEMISTRY AND RADIOGENIC ISOTOPES by LAWRENCE STEPHEN WINTER B.Sc. (Hons.), Memorial University of Newfoundland, 1997 M.Sc., Memorial University of Newfoundland, 2000 A THESIS SUBMITTED IN PARTIAL FULFILLMENT OF THE REQUIREMENTS FOR THE DEGREE OF DOCTOR OF PHILOSOPHY in THE FACULTY OF GRADUATE STUDIES (Geological Science) THE UNIVERSITY OF BRITISH COLUMBIA (Vancouver) APRIL 2008 © Lawrence Stephen Winter, 2008
Abstract The ‘giant’ Tambogrande volcanogenic massive sulphide (VMS) deposits within the Cretaceous Lancones basin of northwestern Perú are some of the largest Cu‐Zn‐Au‐Ag‐bearing massive sulphide deposits known. Limited research has been done on these deposits, hence the ore forming setting in which they developed and the key criteria that permitted such anomalous accumulation of base‐metal sulphides are not understood. Based on field relationships in the host volcanic rocks and U‐Pb geochronology, the deposits formed during the early stages of arc development in the latest Early Cretaceous and were related to an extensional and arc‐rift phase (~105‐100 Ma, phase 1). During this time, bimodal, primitive basalt‐dominant volcanic rocks were erupted in a relatively deep marginal basin. Phase 1 rhyolite is tholeiitic, M‐type, and considered to have formed from relatively high temperature, small batch magmas. The high heat flow and extensional setting extant during the initial stages of arc development were essential components for forming a VMS hydrothermal system. The subsequent phase 2 (~99‐91 Ma) volcanic sequence comprises more evolved mafic rocks and similar, but more depleted, felsic rocks erupted in a relatively shallow marine setting. Phase 2 is interpreted to represent late‐stage arc volcanism during a waning extensional regime and marked the transition to contractional tectonism. The Tambogrande deposits are particularly unusual amongst the ‘giant’ class of VMS deposits in that deposition largely occurred as seafloor mound‐type and not by replacement of existing strata. Paleomorphology of the local depositional setting was defined by seafloor depressions controlled by syn‐volcanic faults and rhyolitic volcanism. The depressions were the main controls on distribution and geometry of the deposits and, due to inherently confined hydrothermal venting, enhanced the efficiency of sulphide deposition. Geochemical and radiogenic isotope data indicate that the rhyolites in the VMS deposits were high temperature partial melts of the juvenile arc crust that had inherited the isotopic signatures of continental crust. Moreover, Pb isotope data suggest the metal budget was sourced almost wholly from mafic volcanic strata. Therefore, unlike the implications of many conventional models, the felsic volcanic rocks at Tambogrande are interpreted to have only played a passive role in VMS formation. ii Table of Contents Abstract ......................................................................................................................................................................... ii Table of Contents ......................................................................................................................................................... iii List of Tables ................................................................................................................................................................. vi List of Figures ............................................................................................................................................................... vii Acknowledgements ..................................................................................................................................................... xii Dedication .................................................................................................................................................................. xiii Co‐Authorship Statement ........................................................................................................................................... xiv Chapter 1. Giant Volcanogenic Massive Sulphide Deposits, Tambogrande, NW Perú 1.1 Introduction ...................................................................................................................... 1 1.2 Background and Approach ................................................................................................ 2 1.3 History ............................................................................................................................... 4 1.4 VMS Deposit Classification and Genetic Models .............................................................. 5 1.5 Controls on ‘Giant’ VMS systems ...................................................................................... 7 1.6 Thesis Objectives............................................................................................................... 9 1.7 Methodology ................................................................................................................... 10 1.7.1 1.7.2 1.7.3 1.7.4 1.7.5 Core Logging .................................................................................................................................... 10 Regional Mapping ............................................................................................................................ 10 Geochronology ................................................................................................................................ 11 Lithogeochemistry ........................................................................................................................... 11 Isotope Chemistry ............................................................................................................................ 12 1.8 Presentation .................................................................................................................... 12 1.9 References ...................................................................................................................... 22 Chapter 2. Volcanic Stratigraphy and Geochronology of the Cretaceous Lancones Basin, Northwestern Perú 2.1 Overview ......................................................................................................................... 27 2.2 Introduction .................................................................................................................... 28 2.3 Tectonic Setting .............................................................................................................. 30 2.4 Regional Geology ............................................................................................................ 31 2.5 Volcanic Stratigraphy ...................................................................................................... 32 2.5.1 2.5.2 2.5.3 2.5.4 Cerro San Lorenzo Formation .......................................................................................................... 35 Cerro El Ereo Formation .................................................................................................................. 37 La Bocana Formation ....................................................................................................................... 38 Lancones Formation ........................................................................................................................ 39 2.6 Structural Geology .......................................................................................................... 39 2.7 Plutonic Rocks ................................................................................................................. 41 2.8 U‐Pb Geochronologic Data ............................................................................................. 42 2.8.1 2.8.2 Volcanic Rocks of the Cerro San Lorenzo Formation ....................................................................... 43 Volcanic Rocks of the Lancones Formation ..................................................................................... 44 2.9 Discussion........................................................................................................................ 47 2.9.1 2.9.2 2.9.3 2.9.4 Depositional Evolution of the Lancones Basin ................................................................................. 47 Timing and Duration of the Volcanic Arc ......................................................................................... 49 Age of Massive Sulphide Deposits ................................................................................................... 50 Comparison of the Lancones Basin to the Western Peruvian Trough ............................................. 51 iii 2.9.5 2.9.6 Tectonic Implications ....................................................................................................................... 52 Inheritance in Zircons and Implications for Basement Rocks .......................................................... 53 2.10 Conclusions ..................................................................................................................... 54 2.11 References ...................................................................................................................... 78 Chapter 3. A Reconstructed Cretaceous Depositional Setting for Giant VMS Deposits at Tambogrande, NW Perú 3.1 Manuscript Status ........................................................................................................... 83 3.2 Abstract ........................................................................................................................... 83 Chapter 4. Volcanic Rock Geochemistry and the Geodynamic Setting of VMS Deposits at Tambogrande, Perú 4.1 Overview ......................................................................................................................... 85 4.2 Introduction .................................................................................................................... 87 4.3 Tectonic Setting .............................................................................................................. 89 4.4 Regional Geology and Volcanic Stratigraphy .................................................................. 91 4.4.1 4.4.2 4.4.3 4.4.4 Cerro San Lorenzo Formation .......................................................................................................... 93 Cerro El Ereo Formation .................................................................................................................. 94 La Bocana Formation ....................................................................................................................... 95 Lancones Formation ........................................................................................................................ 95 4.5 Lithogeochemistry .......................................................................................................... 95 4.5.1 4.5.2 4.5.3 Sampling Procedures and Analytical Methods ................................................................................ 95 Alteration and Element Mobility ..................................................................................................... 96 Geochemical Results ........................................................................................................................ 97 4.5.3.1 4.5.3.2 4.5.3.3 4.5.3.4 Cerro San Lorenzo Formation .............................................................................................................. 97 Cerro El Ereo Formation ....................................................................................................................... 99 La Bocana Formation ........................................................................................................................... 99 Chemostratigraphy of the VMS‐Host sequence ................................................................................. 101 4.6 Discussion...................................................................................................................... 102 4.6.1 4.6.2 4.6.3 Petrochemical Variations in Mafic Volcanic Rocks of the Lancones Basin .................................... 102 Felsic Volcanic Rock Petrochemistry and Association with VMS ................................................... 104 Implications for the Tectonic Setting ............................................................................................. 108 4.7 Summary ....................................................................................................................... 110 4.8 References .................................................................................................................... 135 Chapter 5. Pb‐Sr‐Nd Isotope Systematics of Cretaceous Arc Volcanic Rocks in the Lancones Basin near Tambogrande, Perú – Implications for VMS Deposit Formation 5.1 Overview ....................................................................................................................... 142 5.2 Introduction .................................................................................................................. 144 5.3 Regional Geology and Tectonic Setting ........................................................................ 145 5.4 Volcanic Stratigraphy of the Lancones Basin ................................................................ 148 5.5 Andean Isotopic Framework ......................................................................................... 150 5.6 Pb, Sm‐Nd and Rb‐Sr Isotope Geochemistry ................................................................ 152 5.7 Analytical Methods ....................................................................................................... 153 5.7.1 5.7.2 Pb Isotope Analysis, Mineral Separates ......................................................................................... 153 Pb, Rb‐Sr, Sm‐Nd Isotope Analysis, Whole Rock Samples ............................................................. 154 5.8 Results ........................................................................................................................... 155 5.8.1 5.8.2 Volcanic Rocks ............................................................................................................................... 155 Massive Sulphide Deposits ............................................................................................................ 158 5.9 Discussion...................................................................................................................... 159 iv 5.9.1 5.9.2 Volcanic Rock Petrogenesis ........................................................................................................... 159 Possible Linkages Between Petrogenesis of the Felsic Volcanic Suite and VMS Formation .......... 163 5.10 Summary and Conclusions ............................................................................................ 166 5.11 References .................................................................................................................... 187 Chapter 6. Summary, Discussion and Unresolved Questions 6.1 Summary and Main Conclusions ................................................................................... 192 6.2 Discussion and Ideas ..................................................................................................... 198 6.2.1 6.2.2 Timing and Tectonic Setting .......................................................................................................... 198 Where’s the Intrusion? .................................................................................................................. 199 6.3 Outstanding Issues and Directions for Future Research .............................................. 200 6.3.1 6.3.2 6.3.3 At Tambogrande and in the Lancones Basin ................................................................................. 200 In the Western Peruvian Trough ................................................................................................... 201 Globally .......................................................................................................................................... 203 6.4 References .................................................................................................................... 205 A.1 A.2 A.3 A.4 Appendix A. U‐Pb Zircon Sample Preparation, Analysis and Additional Data Methodology ................................................................................................................. 210 Additional U‐Pb Zircon Geochronologic Data ............................................................... 217 Results ........................................................................................................................... 217 References .................................................................................................................... 225 Appendix B. Ar‐Ar Geochronologic Data B.1 Methodology ................................................................................................................. 226 B.2 Results ........................................................................................................................... 227 B.3 References .................................................................................................................... 236 Appendix C. Lithogeochemistry C.1 Analytical Methods, Precision, and Accuracy ............................................................... 237 C.2 References .................................................................................................................... 260 v List of Tables Table 1.1. Individual deposit data from number drill holes, tonnage and grade (Manhattan Minerals, 2002). ........ 21 Table 2.1. Summary of location and description data for samples analyzed in this study for U‐Pb zircon dating. .... 58 Table 4.1 (Following page). Average lithogeochemical values and 2 σ errors for volcanic rocks of the Lancones basin, based on volcanic rock type and formation. Complete data are listed in Appendix C. ........................ 128 Table 4.2. Summary of average selected trace element ratios for volcanic rocks of the Lancones basin based on volcanic rock type and formation. ................................................................................................................... 133 Table 5.1. Sample location data, approximate age and rock descriptions. All samples are from diamond drill core except for LW002. Coordinates are in map projection WGS 84, UTM Zone 17 Southern Hemisphere. ........ 184 Table 5.2. Pb, Nd and Sr isotope data from volcanic rocks associated with VMS deposits at Tambogrande. ‘Initial’ isotope values are calculated to 100 Ma. ........................................................................................................ 185 Table 5.3. Pb isotope compositions of Tambogrande ore deposits and post‐mineralization intrusive phases in the Lancones basin. ................................................................................................................................................ 186 Table A1. U‐Pb zircon analytical data obtained using ID‐TIMS method. .................................................................. 211 Table A2. U‐Pb zircon analytical data obtained using the SHRIMP‐RG method. ...................................................... 212 Table A3. Description of additional rock samples for U‐Pb zircon analysis (not included in the Chapters). ............. 218 Table A4. U‐Pb zircon data from SHRIMP‐RG analysis. ............................................................................................ 219 Table B1. 40Ar‐39Ar rock sample description and location data. Eastings and northings are UTM Zone 17, Southern Hemisphere (WGS84). ..................................................................................................................................... 228 Table B2. 40Ar‐39Ar age data for plutonic and volcanic rock samples from the Lancones basin. Neutron flux monitors: 24.36 Ma MAC‐83 biotite (Sandeman et al. 1999); 28.02 Ma FCs (Renne et al., 1998). Isotope production ratios: (40Ar/39Ar)K=0.0302, (37Ar/39Ar)Ca=1416.4306, (36Ar/39Ar)Ca=0.3952, Ca/K=1.83(37ArCa/39ArK). ................................................................................................................................. 229 Table C1. Whole rock geochemical analyses. Major oxides are from Bondar‐Clegg Laboratories. Trace elements are from Memorial University. Abbreviations: CSLF = Cerro San Lorenzo Formation; CEEF = Cerro El Ereo Formation; LBF = La Bocana Formation. D = dacite; A = andesite; B = basalt; R = rhyolite; RD = rhyolite dyke (post mineralization); bx = breccia. .................................................................................................................. 239 Table C2. Whole rock geochemical analyses. Major oxides are from Bondar‐Clegg Laboratories. Trace elements are from Memorial University. ........................................................................................................................ 246 Table C3. Memorial University analyses of in‐house standards run with samples from this study. ........................ 256 Table C4. ALS Chemex analyses of MDRU standards run with samples from this study. ......................................... 257 vi List of Figures Figure 1.1. Location maps and simplified geology for the study area. The locations of VMS deposits (TG1, TG3, and B5) in the Tambogrande area are also shown and field area of this study outlined. Geology modified after Jaillard et al. (1999) and Tegart et al. (2000). .................................................................................................... 15 Figure 1.2. Schematic model for the formation of VMS deposits (from Franklin et al., 2005). .................................. 16 Figure 1.3. Schematic section and model of a typical volcanogenic massive sulphide deposit from modern mid‐
ocean ridge settings; after Herzig and Hannington (1995). ............................................................................... 17 Figure 1.4. Histogram of ages for global bimodal‐mafic type VMS deposits (n=327); data from Franklin et al. (2005).
........................................................................................................................................................................... 18 Figure 1.5. Metals versus size of the deposit (tonnes) for global VMS deposits of the bimodal‐mafic class (n=326; data from Franklin et al., 2005). Tambogrande deposits are labeled. KC = Kidd Creek deposit. A. Copper and B. Zinc ................................................................................................................................................................. 19 Figure 1.6. Gold grade (grams/tonne) versus size of the deposit (tonnes) for global VMS deposits of the bimodal‐
mafic class (n=326; data from Franklin et al., 2005). Tambogrande deposits are labeled. KC = Kidd Creek deposit. .............................................................................................................................................................. 20 Figure 2.1. Morphostructural units of the Peruvian Andes (modified after Benavides‐Cáceres, 1999). Cretaceous marginal basins ‐ Lancones (LB), Huarmey (HB) and Cañete (CB) basins ‐ are superimposed. Also shown are the locations of VMS deposits and prospects (circles) (data from Steinmüller et al., 2000). ............................ 56 Figure 2.2. A. Location map for the Tambogrande project; B. regional map showing major tectonostratigraphic units of coastal northwestern Perú. The locations of VMS deposits (TG1, TG3, and B5) in the Tambogrande area are also shown and field area of this study outlined. Modified after Jaillard et al. (1999) and Tegart et al. (2000). ................................................................................................................................................................ 57 Figure 2.3 (on the following page). Schematic paleogeographic model of the development of Perú‐Ecuador segment of the western margin of South America (SA) from the Jurassic to present using data from Mourier et al. (1988), Mitouard et al. (1990), Litherland et al. (1994), Aspden et al. (1995), Noble et al. (1997), Arculus et al. (1999), Benavides‐Cáceres (1999), Jaillard et al. (2000), Bosch et al. (2002), and Polliand et al. (2005). A. Jurassic to earliest Early Cretaceous: ~SE‐directed convergence of the proto‐Farallon‐Caribbean ocean plate with continental SA; subduction occurs along the Ecuadorian segment, whereas the Peruvian NNW‐trending margin is a sinistral transform; Amotape terrane is a micro‐continent approaching SA; B) change in convergence direction from SE to ~NE; accretion of the Amotape terrane, notably along the Peruvian segment; dextral faulting of Amotape terrane and clockwise rotation of blocks; ocean‐continental plate boundary ‘jumps’ toward the west; during this period the NW‐trending Peruvian margin becomes a subduction zone whereas the Ecuadorian NE‐trending margin becomes a dextral transform. C) trench ‘roll‐
back’ occurring along Peruvian margin and extension in overriding SA plate; the Lancones basin and Western Peruvian Trough open up along a margin parallel rift and result in the deposition of Cretaceous sedimentary and arc volcanic rocks; continued dextral displacement of Amotape terrane; D) termination of marginal rifting, accretion of ocean plateau ‘Pallatanga’ terrane in Ecuador, and deformation of Andean terranes; formation of Macuchi island arc near margin to be accreted to Ecuadorian segment by Early Oligocene; E) modern day tectonostratigraphic model; compressive tectonic regime; E‐directed convergence. .................. 60 Figure 2.4. Regional geologic map for the portion of the Lancones basin reviewed in this study; modified from Reyes and Caldas (1987) and from mapping during this study. The location of samples for U‐Pb zircon geochronologic studies are shown and labeled by sample name. Lines A‐A’ and B‐B’ show the trace of sections in figures 2.5 a‐b. ................................................................................................................................. 62 Figure 2.5. A. Regional geological cross section A‐A’ through the southern region of the map area. Looking northeast. TG1 and TG3 massive sulphide deposits projected from the south. B. Regional geological cross section B‐B’ through the northern region of the map area. Looking northeast. Legend as per figure 2.4. See map in Figure 2.4 for trace of sections. ............................................................................................................. 63 Figure 2.6. Schematic stratigraphic column of the eastern portion of the Lancones basin. Inset section shows the Tambogrande area in more detail. .................................................................................................................... 64 Figure 2.7 (following page). Field and drill core photographs of mafic rocks from the Cerro San Lorenzo Formation: A. Feldspar porphyritic and amygdaloidal basalt. B. Drillcore from the B5 area, aphyric basalt with autobrecciated margin and close up of breccia C. Illustrates the highly vesicular, scoria‐like clasts; note the small fragments of bubble‐wall shards. D. Section through basaltic pillow lavas at Rio Quiroz; pillows are up to 1 m wide; individual pillow flows are up to 10’s of metres thick. E. Basaltic pillow lavas displaying well vii developed concentric flow foliations; this specimen is partly broken along radial fractures. F. Pillow basalt unit (P) overlain by mass flow (MF) deposit of basaltic pillow lava and autobreccia clasts. A mafic dyke (Dk) cuts both units and would have probably supplied lava to another cycle of pillow lavas and breccias. G. Medium to thick bedded basaltic volcaniclastic deposits ranging from sand‐ to boulder‐size; note the reverse sorting of the thicker (~1m) basal unit (see arrow) possibly indicative of a massive flow. H. In‐situ autoclastic (hyaloclastic) breccia. Note the jigsaw‐fit textures of the clasts. Breccia are gradational into massive lavas. Drill core, B5. I. In‐situ autoclastic breccias from drillcore, TG3. Bulbous‐shaped clast with diffuse margins in a dark green chlorite matrix; margins of clasts display fine (sub mm) chloritic amygdules. This breccia grades into massive lava. ............................................................................................................................................... 65 Figure 2.8. Drill core photographs of intermediate and felsic rocks from the Cerro San Lorenzo Formation: A. Massive feldspar porphyritic rhyolite. Scale units are mm. TG1 area. B. Flow‐banded rhyolite autobreccia; these breccias typically grade into lavas. Textures partly masked by quartz and sericite alteration. TG3 area. C. Rhyolitic, unsorted, clast‐supported, volcaniclastic rock with pebble‐size clasts and massive sulphide fragments (near the TG3 deposit). D. Green‐grey, feldspar porphyritic dacite with large flow‐foliated chlorite‐quartz ‘pipe’ amygdules. Hanging wall to TG3 deposit. ...................................................................... 67 Figure 2.9. Field photographs of mafic rocks from the Cerro El Ereo Formation: A. Typical porphyritic textures of the Cerro El Ereo Formation porphyritic basalt. Sample contains ~20% feldspar phenocrysts to >1 cm in an aphanitic matrix. Non‐amygdaloidal. Subvolcanic or thick flow facies. B. Bleached‐looking, boulder size, subround clasts (C) of basalt feldspar porphyry in a fine matrix of dark grey feldspar porphyritic material (M). Clasts show in‐situ breccia textures (jigsaw‐fit) attributed to progressive fragmentation of blocks during transport. C. Unsorted, non‐stratified basaltic cobble‐ to pebble‐sized lithic and feldspar crystal‐bearing volcaniclastic rock. The sample contains an equal proportion of aphyric to weakly feldspar porphyritic (W) clasts and coarse feldspar (C) porphyry clasts. Amygdaloidal clasts (A) are present but are generally not common. Clast margins often are not easily discernible. D. Thin‐ to thick‐bedded feldspar crystal to ash‐
sized tuff; reworked facies at top of formation. ................................................................................................ 68 Figure 2.10. Field photographs of mafic rocks from the La Bocana Formation: A. Moderately west‐dipping, thick massive basaltic‐andesite flows. Felsic stocks and dykes cut perpendicular to bedding. B. Basaltic andesite dykes with strongly flow foliation defined by flattened and large vesicules (silica amygdules up to 30 cm). C. Polylithic, basaltic‐andesite dominated, mass flow deposit. Not the fractures and in‐situ fragmentation of the clasts due to mass transport (see arrows). D. Cracked and brecciated outer crust of basaltic andesite lava flow and interstitial hyaloclastite resulting from quenching of exposed lava. E. Mafic lobes (M) injected into felsic quartz‐crystal tuffs (T). Tuffs show ‘soft‐sediment’ deformation textures and mafic flows show columnar jointing indicating tuffs were non‐welded/non‐lithified during deposition of mafic flows. F. Lithic and quartz‐feldspar crystal rhyolitic tuff. Colouration of the domains are a result of secondary recrystallization to quartzo‐feldspathic (light) and chloritic (+clay) assemblages due to devitrification of glass component. ........................................................................................................................................................ 69 Figure 2.10 (continued). G. Rhyolitic quartz crystal‐rich and lithic tuff. H. Coarse, boulder breccia with chaotic, unsorted, subround (pillow?) clasts. Talus breccia. I. Medium bedded, well sorted and locally cross bedded (arrow), pebble‐ to sand‐sized, mafic‐dominated volcaniclastic rocks. ............................................................. 70 Figure 2.11. A. thin bedded arenaceous sequence from the Lancones Formation; massive unit at top of outcrop is a diorite sill. B. thin bedded limestones and limey‐arenites. ............................................................................... 71 Figure 2.12. Schematic stratigraphic column of the eastern portion of the Lancones basin. Legend as per Figure 2.6. Inset section shows the Tambogrande VMS section in more detail. Age data from this study are shown in their relative stratigraphic positions. Ages of plutonic rocks provided herein (Appendices A, B). ............... 72 Figure 2.13. (following page) 238U/206Pb versus 207Pb/206Pb Tera‐Wasserburg plots (Tera and Wasserburg, 1972) for various volcanic rock samples from the Cerro San Lorenzo and La Bocana Formations. Error ellipses are 2σ. Dashed lines indicate data points omitted versus solid lines/grey ellipses for data included in the age calculation. Inset figures show box plots for all sample points for 207Pb‐corrected 206Pb*/238U data with error bars at 2σ. Open boxes are omitted whereas solid boxes were included in the age calculation. Ages given are 206Pb/238U with 2σ uncertainties. ................................................................................................................. 73 Figure 2.14. 207Pb/235U versus 206Pb/238U U‐Pb concordia plots for various volcanic rock samples from the Cerro San Lorenzo and La Bocana Formations. ........................................................................................................... 75 Figure 2.15. Schematic stratigraphy and U‐Pb zircon ages that constrain the volcanic formations in the Lancones basin. .................................................................................................................................................................. 76 viii Figure 2.16. Comparison of schematic volcanic stratigraphy of the Lancones Basin and Western Peruvian Trough (modified from Myers, 1974; Offler et al., 1980; Cobbing et al., 1981) with emphasis on age correlation. Legend as per Figure 2.6. ................................................................................................................................... 77 Figure 4.1. Morphostructural units of the Peruvian Andes (modified after Benavides‐Cáceres, 1999). Cretaceous marginal basins ‐ Lancones (LB), Huarmey (HB) and Cañete (CB) basins ‐ are superimposed. Also shown are the locations of VMS deposits and prospects (circles) (data from Steinmüller et al., 2000). .......................... 111 Figure 4.2. A. Location map for the Tambogrande project; B. regional map showing major tectonostratigraphic units of coastal northwestern Perú. The locations of VMS deposits (TG1, TG3, and B5) in the Tambogrande area are also shown and field area of this study outlined (see Fig. 4.3 for a detailed map). Modified after Jaillard et al. (1999), Tegart et al. (2000). ........................................................................................................ 112 Figure 4.3 – Location map and simplified cross sections along the Peruvian continental margin based on gravity modeling and seismic data from Couch et al. (1981) and Jones (1981). ......................................................... 113 Figure 4.4 (following page). Regional geologic map for the Tambogrande area of the Lancones basin reviewed in this study. The location of VMS deposits TG1, TG3, and B5, as well as geochemical sampling locations are shown. Map projection is WGS 84, Zone 17S. Map is from this study. ......................................................... 114 Figure 4.5. Schematic stratigraphic column of the volcanic arc sequence of the Lancones basin. Inset section is a more detailed schematic section of the VMS‐bearing sequence at Tambogrande. ........................................ 116 Figure 4.6. A. Silica vs. total alkalies classification scheme of Le Bas et al. (1986). B. Nb/Y versus Zr/TiO2 plot of Winchester and Floyd (1977). C. AFM plot (Irvine and Baragar, 1971). ......................................................... 117 Figure 4.7. Bivariate plots of basalt from the Cerro San Lorenzo, Cerro El Ereo and La Bocana formations. A. SiO2 versus MgO, B. Ni versus MgO, C. Cr versus MgO, and D. V versus MgO. The vertical dashed line at 5.5 w% MgO emphasizes the division in the data. ....................................................................................................... 118 Figure 4.8. Basalt discrimination diagrams. A. Th‐Zr‐Nb plot (Wood, 1980). B. Zr‐Nb‐Y plot (Meschede, 1986); all samples illustrate relatively low Nb but variable Th, Zr, and Y values and are defined as arc basalt. C. Zr‐Ti‐Y plot (Pearce and Cann, 1973). D. Ti‐V plot (Shervais, 1982). ......................................................................... 119 Figure 4.9. Primitive mantle‐ normalized extended trace element diagrams for mafic intermediate rocks from the various formations. A. Cerro San Lorenzo Formation basalt. B. Cerro San Lorenzo Formation basaltic‐
andesite. C. Cerro El Ereo Formation basalt. D. Cerro El Ereo Formation basaltic‐andesite. D. La Bocana Formation basalt. E. La Bocana Formation basaltic‐andesite. Element order and normalizing values follow Sun and McDonough (1989). ........................................................................................................................... 120 Figure 4.10. Chondrite‐normalized (using values from Sun and McDonough, 1989) HFSE values for basalt from the Cerro San Lorenzo, Cerro El Ereo, and La Bocana formations. A. Yb versus La/Yb. B. Y versus Zr/Y. ........... 121 Figure 4.11. Felsic volcanic discrimination diagrams. A. Ga/Al versus Zr after Whalen et al. (1987). B. Y versus Nb (Pearce et al., 1984). All samples plot within the I‐ to M‐type field. .............................................................. 122 Figure 4.12. Primitive mantle‐normalized extended trace element diagrams show broadly similar patterns for felsic volcanic rocks of the Lancones basin. A. Cerro San Lorenzo Formation. B. La Bocana Formation. Inset diagrams are rare earth elements only and are normalized to chondrite values. Element order and normalizing values follow Sun and McDonough (1989). ................................................................................. 123 Figure 4.13 (following page). Multi‐element bivariate plots for basalt, dacite, and rhyolite in the vicinity of VMS deposits at Tambogrande. A. Zr/TiO2 vs. Nb/Y (Winchester and Floyd, 1977). B. Na2O+K2O vs. SiO2 (Le Bas et al., 1986). C. Fe2O3+MgO vs. SiO2. D. TiO2 vs. Zr. E. P2O5 vs. Zr. F. V vs. Zr/TiO2. G. Hf vs. Nb. H. Zr/Y vs. Y. I. La/Yb vs Yb (chondrite‐normalized values using Sun and McDonough (1989)). J. Th vs. Zr. ......................... 123 Figure 4.14. Ta/Yb versus Th/Yb plot after Pearce (1983). Only Cerro San Lorenzo Formation and El Ereo Formation basalt shown due to data limits for La Bocana Formation. D = depleted mantle. E = enriched mantle. ............................................................................................................................................................. 125 Figure 4.15. Schematic two‐stage tectonic model for arc magmatism in the Lancones basin. The model proposes a shift from phase 1, extensional tectonics with a steeply dipping subduction zone to phase 2, waning extension and shallower subduction. Depletion of the mantle‐wedge explains the relatively HFSE‐ and REE‐
depleted mafic volcanic rocks in the phase 2. The thickened crust due to waning extension forces more fractionation of mafic intrusions. The partial melting of the arc crust causes phase 2 felsic volcanic rocks to yield lower HFSE contents. ............................................................................................................................... 126 Figure 4. 16. Felsic volcanic rock discrimination diagrams (after Lesher et al., 1986; Barrie et al., 1993; Lentz, 1998; Hart et al., 2004). A. (La/Yb)N versus YbN (Chondrite‐normalized using values from Sun and McDonough (1989)). B. Zr/Y versus Y. ................................................................................................................................ 127 ix Figure 5.1. Morphostructural units of the Peruvian Andes (modified after Benavides‐Cáceres, 1999). Cretaceous marginal basins ‐ Lancones (LB), Huarmey (HB) and Cañete (CB) basins ‐ are superimposed. Also shown are the locations of VMS deposits and prospects (circles) (data from Steinmüller et al., 2000). .......................... 169 Figure 5.2. A. Location map for the Tambogrande project B. Regional map showing major tectonostratigraphic units of coastal northwestern Perú. The locations of VMS deposits (TG1, TG3, and B5) in the Tambogrande area are also shown and field area of this study outlined (see Fig. 5.3 for a detailed map). Modified after Jaillard et al. (1999), Tegart et al. (2000). ........................................................................................................ 170 Figure 5.3. Location map and simplified cross sections along the Peruvian continental margin based on gravity modeling and seismic data from Couch et al. (1981) and Jones (1981). ......................................................... 171 Figure 5.4 (next page). Regional geologic map for the Tambogrande area of the Lancones Basin. The location of VMS deposits TG1, TG3, and B5, where the bulk of the isotope samples were collected are shown Other individual samples from the region are labeled. Map projection is WGS 84 (World Geodetic System), UTM Zone 17 Southern Hemisphere. ....................................................................................................................... 172 Figure 5.5. Schematic stratigraphic column of the volcanic arc sequence of the Lancones basin. Inset section shows a more detailed schematic section of the VMS‐bearing sequence at Tambogrande. .......................... 174 Figure 5.6. Map depicting Pb isotope provinces of the Andes (modified after Macfarlane et al., 1990; Tosdal et al, 1999). The location of Tambogrande and other VMS deposits within the Cretaceous marginal basins of Perú are shown. ....................................................................................................................................................... 175 Figure 5.7. Thorogenic (A) and Uranogenic (B) Pb isotope diagrams for data from the Lancones basin and fields for Pb isotope provinces of the Andes (after Macfarlane et al., 1990). All data points are from this study. Symbols for rocks and sulphide samples and fields for Pb provinces are given in inset boxes in B. S & K = Stacey and Kramers (1975) growth curve. ....................................................................................................... 176 Figure 5.8. Schematic east‐west cross section through the Lancones basin showing the main tectonic units and the spatial distribution of units sampled for isotopic analysis (including this study and data available from the literature; see references in the text). ............................................................................................................. 177 Figure 5.9. Thorogenic (A) and Uranogenic (B) Pb isotope diagrams for data from the Lancones Basin and fields for Pb isotope signatures of various tectonic units of the northern Andes or Perú and Ecuador. Data for ’Cretaceous platform sedimentary rocks’ and ‘continental crust (Olmos Complex)’ from the central Andes, Perú, after Macfarlane et al. (1990) and Macfarlane (1999). Data for ‘Jurassic‐Cretaceous metasedimentary rocks’ and ‘Jurassic‐Cretaceous MORB’, from Ecuador, after Bosch et al. (2002). Data for ‘Coastal Batholith’ from Mukasa (1986) and this study. Data for ‘East Pacific MORB’, East Pacific Rise, from Sun (1980). All data points are from this study. Symbols for rocks and sulphide samples and fields for Pb provinces are given in inset boxes in B. ............................................................................................................................................... 178 Figure 5.10. Rb‐Sr isotope plots for volcanic rocks of the Lancones Basin. A. Sr versus 87Sr/86Srί. B. 87Sr/86Sr versus 87
Rb/86Sr. The line shown is defined by the felsic volcanic rocks only (n=6) and yields an errorchron of ~65 Ma. ................................................................................................................................................................... 179 Figure 5.11 ‐ 143Nd/144Nd versus 87Sr/86Sr for volcanic rocks of the Lancones basin. A. Compared to fields for regional geologic units. B. Enlargement of data shown in A. Fields are given in the legend in A and symbols for both plots are shown in B. Data sources as per Fig. 5.9. ........................................................................... 180 Figure 5.12. Geochemical discrimination diagrams for volcanic rocks from the Lancones Basin. A. εNd versus Th/Yb. B. 206Pb/204Pb versus Th/Yb. ............................................................................................................................ 181 Figure 5.13. Uranogenic Pb isotope diagram showing isotope compositions for mafic and felsic volcanic rocks at Tambogrande, as well as ore mineral isotope compositions. Volcanic rock data is conceptually time‐adjusted to 100 Ma based on Stacey and Kramers (1975) crustal growth curve with µ=9.85. Mixing lines are shown and labeled. ..................................................................................................................................................... 182 Figure 5.14 ‐ Conceptual magmatic‐hydrothermal model relating petrogenesis of bimodal mafic‐felsic volcanic rocks at Tambogrande to hydrothermal system that formed VMS deposits (adopted from the petrogenetic model of Hart et al. (2004) for FII‐FIII felsic volcanic rocks). Depth of magma chamber is suggested by geochemical data which indicate partial melting or fractionation with amphibole ± pyroxene and plagioclase at shallow crustal depths. Isotope data support partial melting models as felsic volcanic rocks are substantially different isotopically as compared to basalts, and, yield more heterogeneous isotopic results. Hydrothermal leaching of metals, which is limited by the depth of fracture permeability, did not penetrate 206
204
crustal rocks that were responsible for the unique isotope values (high Pb/ Pb) in the VMS‐associated felsic rocks. ....................................................................................................................................................... 183 Figure A1. SEM Cathodoluminesence images of zircons showing location of spot analyses for SHRIMP‐RG data. . 216 x Figure A2. Regional geologic map of northwestern Perú showing the location of U‐Pb zircon geochronology samples. ........................................................................................................................................................... 222 Figure A3. Box plot for all sample points for sample LW‐30 illustrating 207Pb‐corrected 206Pb*/238U data with error bars at 2σ. Open boxes are omitted whereas solid boxes were included in the age calculation. Ages given are 206Pb/238U with 2σ uncertainties. ............................................................................................................... 223 Figure A4. Box plot for all sample points for sample LW‐31 illustrating 207Pb‐corrected 206Pb*/238U data with error bars at 2σ. Open boxes are omitted whereas solid boxes were included in the age calculation. Ages given are 206Pb/238U with 2σ uncertainties. ............................................................................................................... 223 Figure A5. Box plot for all sample points for sample 3763 illustrating 207Pb‐corrected 206Pb*/238U data with error bars at 2σ. Open boxes are omitted whereas solid boxes were included in the age calculation (and are noted by those encircled with the dashed line). Ages given are 206Pb/238U with 2σ uncertainties. .......................... 224 Figure A6. Box plot for all sample points for sample 1601 illustrating 207Pb‐corrected 206Pb*/238U data with error bars at 2σ. No age was determined for this sample, though the circled data points represent possible igneous age of the sample. Ages given are 206Pb/238U with 2σ uncertainties. ................................................ 224 Figure B1 ‐ Geology map of study area showing the location of Ar‐Ar samples. Map projection is UTM Zone 17, Southern Hemisphere (WGS84). ...................................................................................................................... 232 Figure B2 – Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐06 (granodiorite). ........ 233 Figure B3 ‐ Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐07 (rhyolitic volcaniclastic)
......................................................................................................................................................................... 233 Figure B4 – Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐36 (diorite). .................. 234 Figure B5 ‐ Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐88 (hornblende porphyritic dyke). ............................................................................................................................................................... 234 Figure B6 ‐ Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐85 (hornblende granite).
......................................................................................................................................................................... 235 Figure B7 ‐ Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐61 (hornblende‐plagioclase porphyritic dyke). ............................................................................................................................................. 235 Figure C8. (following page). Primitive mantle‐normalized trace element diagrams for repeat analyses of in‐house and internal reference material conducted during this study. A. ICP‐MS data (Memorial Univ.) for standards MRG‐1 and BR‐688 compared to the detection limit. B. Average analyses for the reference materials from this study in (A) compared to given values from previous analysis by Memorial Univ. of the material. C. ICP‐
AES data (ALS Chemex) for MDRU reference samples BAS‐1 and P‐1 as compared to the detection limit for the technique. D. Average values for the reference materials from this study from in (C) compared to data from Piercey (2001). ........................................................................................................................................ 258 xi Acknowledgements This project began at the Roundup conference in Vancouver in January, 1999, when Steve Piercey introduced me to Jim Mortensen. I thank both of them for getting me started on this journey. Dick Tosdal agreed to take me on a year later, though I doubt he thought it might be 2008 before the final product by his first doctoral student would be delivered! I certainly didn’t think it would take 8 years. However, Dick has shown incredible patience and tolerance for which I am very grateful. Without your constant support and encouragement this project would not have been completed. Few people have worked on and studied VMS systems as much as Jim Franklin. I recall as an undergrad reading many papers on VMS deposits by J.M. Franklin. I thank Jim for sharing his knowledge and for working with us on this project, for many great stories, and for providing excellent advice on both academic and a few not‐so‐academic issues. Peter Tegart, who spearheaded Manhattan Minerals, initiated this project and deserves much credit for the recognition of the economic potential of Tambogrande district. Peter was instrumental in getting the company to fund this project funded and also provide logistical support. I thank him for giving me the opportunity to work at Tambo. Additional financial support for this project came from a NSERC Collaborative Research and Develop grant, a Hugh E. Mckinstry Grant (Society of Economic Geologists Foundation), the Egil H. Lorntzsen and Thomas and Marguerite MacKay Memorial scholarships, UBC. The Manhattan Minerals exploration and development team in Peru made working there an unforgettable experience while contributing significantly to the ideas that have become part of this thesis. I am particularly grateful to Andy Carstensen, Cristian Soux, Gord Allen, and Brian Thurston, Kosta and Sefika Lesnikov, and Arturo Cordova for many great discussions of these deposits, as well as for great memories of working/living in Peru. Allan San Martin also helped out significantly with data management. Miguel Jimenez provided excellent field assistance. EOS and MDRU. The geochronology and isotope work for this thesis was supported immensely by many folks at UBC who graciously provided their time and attention to turning my rocks into data, including Janet Gabites (Pb isotopes) Rich Friedman (U‐Pb), Tom Ullrich (Ar). Thanks to Dominique Weis for running Pb‐Sr‐
Nd work. Jim Mortensen kindly allowed me to destroy several of the ceramic grinding plates and worked hard to help squeeze some dates out these rocks. Claire Chamberlain skillfully ran the SHRIMP work. Karie Smith and Arne Toma are thanked for keeping my admin and technical life in order. Steve Piercey and Derek Wilton provided excellent informal reviews of this thesis and their support is most appreciated. Kelly Russell, Malcolm Scoble, and Jan Peter (GSC) served on the examination committee and provided helpful and constructive criticism. Many of my fellow grad students at UBC are thanked for the good discussions and excellent distractions. Thanks especially to Diego (and Diana) Charchaflie, Kathy Dilworth, Scott Heffernan, Steve Israel, Nancy MacDonald, Piercey, Steve Quane, and Dave Smithson. Simon Haynes made living on 16th Ave. a most memorable time in my (our) lives. Geoff Bradshaw was a great friend and officemate whom I shall forever remember. Peace. A big thanks to Altius for allowing me the time and space to get this done, and especially the staff who pulled extra weight for me. A heartfelt thanks to Ken Hickey and Frankie Goodwin for all your help and for being there for us when I (we) really needed a hand. My family has always been extremely supportive of what I have decided to do in life and I am indebted to my parents and sister for their encouragement and love. Nobody has been more supportive or understanding during this undertaking than my wife, Angie. Thank you for your incredible patience, constant support and reassurances when I was least optimistic that this would be completed. You are simply amazing. You, Dylan, and Teagan were my inspiration to finish this venture. Thanks for your love. Vancouver 27 Jan 08. xii Dedication To my Parents
xiii Co‐Authorship Statement This dissertation comprises the sum of the author’s research and has applications to both mineral deposits exploration and academic studies. It comprises a collection of papers that were written for the purpose of publication in economic geology journals and in many cases benefited from the collaboration with various industry and university geologists. There are many geologists who worked in the employment of the supporting mining company during the exploration and development program and who contributed indirectly to the development of this thesis. Notwithstanding the significant contribution of these workers, and except in circumstances where such acknowledgement is provided, the ideas presented herein are entirely those of the author. Unless otherwise indicated, all maps, sections, level plans, 3D perspectives, and all other plots and figures were constructed by the author using data generated by the author. The co‐authors helped refine the concepts of this paper and contributed editorially. xiv Chapter 1. Giant Volcanogenic Massive Sulphide Deposits, Tambogrande, NW Perú 1.1 Introduction Giant ore deposits represent a small number of examples within any particular deposit type, though they commonly contain a major portion of the metal budget within the deposit type (Singer, 1995). Explorationists seek giant ore deposits because of the robust economic value and the stability and longevity of mining that such large tonnage deposits provide. Studies that attempt to understand the key criteria in the genesis of various giant ore deposit types (e.g., Clark, 1993, 1995; Goodfellow and Zierenberg, 1999; Gibson et al., 2000; Hedenquist et al., 2000) typically conclude that such deposits exist because all formational processes operated under optimal conditions and/or that special ore‐forming circumstances were required. Despite years of searching for the key criteria other than direct detection methods, there are limited tools available to assist the explorer in designing effective exploration programs for giant deposit types. Nonetheless, it is clear that a better understanding of the geological processes and conditions essential to the formation of giant deposits will be required to improve exploration methodologies. The Tambogrande volcanogenic massive sulphide (VMS) deposits of northwestern Perú (Fig. 1.1) constitute a remarkable VMS district due to the presence of multiple abnormally large tonnage Cu‐Zn‐Au‐Ag‐bearing massive sulphide deposits. In contrast, most VMS camps contain log‐normally distributed deposit sizes (e.g., 1 large, 2‐3 middle sizes, and 5‐6 small deposits; Sangster, 1972; Franklin et al., 2005), such as at Noranda (Gibson and Watkinson, 1990) or Flin Flon (Syme and Bailes, 1993). The presence of more than one giant massive sulphide within a single cluster is unusual and few regions are known to exhibit this feature (e.g., Urals; Chapter 1 Page 1 Herrington et al., 2005). The three deposits in the Tambogrande district define an unusual resource of massive sulphide mineralization. There is only an indirect understanding of the geological attributes of a district that enabled the formation of giant deposits, and even fewer constraints on the geological processes which permit the generation of multiple giant deposits. As many of these deposits have not been developed to the mining stage, little research has been done to elucidate understanding of the potential world‐class district at Tambogrande. Therefore, the theme of this thesis is to evaluate the tectonic, volcanic, and depositional controls on the formation of the series of giant VMS deposits at Tambogrande. Despite the fact VMS deposits are a relatively well understood hydrothermal metal deposit‐type, a better comprehension of the geology of well‐preserved examples at Tambogrande will assist in refining VMS deposit models and aid in the understanding of the genesis of the largest hydrothermal ore deposits. 1.2 Background and Approach Continental margin arc sequences are known from Ecuador in the north to Tierra del Fuego in southernmost Chile and Argentina (Dalziel, 1981; Atherton et al., 1983; Vergara et al., 1995; Jaillard et al., 1996; Hanson and Wilson, 1991) and represent vestiges of a Mesozoic arc‐
related rift system (i.e., intra‐arc or back‐arc) that developed at the leading edge of western continental South America. A variety of ore deposit types occur throughout the rift sequences, including Chilean manto, skarn, Fe‐oxide Cu‐Au, porphyry Cu‐Au, and VMS deposits, though few VMS deposits other than those at Tambogrande are economically significant. The tectonic and magmatic processes that formed the Lancones basin, the northernmost of the Mesozoic arc‐rift sequences, underpin the processes that enabled the formation of these unusually large deposits. Understanding the igneous‐volcanic and tectonic evolution of the Lancones basin Chapter 1 Page 2 thus provides a basis for understanding the unique genetic attributes that enabled the formation of the Tambogrande deposits. The study area in northwestern Perú is situated within an enigmatic position along the Andes orogen known as the Huancabamba deflection (Mourier et al., 1988; Mitouard et al., 1990). This major oroclinal bend marks an abrupt transition from the north‐northwest trending Peruvian Andes to the north‐northeast trending Ecuadorian Andes. Though the tectonic history of the orocline is not well understood. The Ecuadorian and Peruvian segments also record somewhat different accretionary and magmatic histories (Jaillard et al., 1999, 2000; Benavides‐
Cáceres, 1999) adding to the significance of this transitional zone. A better understanding of the development of the Lancones basin is integral to the development of tectonic models of the Huancabamba deflection and Mesozoic Andean orogenesis. To summarize, this thesis adds to the theoretical and applied knowledge base on giant VMS deposits and to the genesis of VMS deposits by documenting the regional tectonomagmatic setting and also the detailed local volcanological, structural and stratigraphic controls on VMS formation at Tambogrande. The pristine state of preservation of the deposits, well constrained volcanic stratigraphic sections, and good understanding of the tectonic history of the region provide excellent controls to establish the geologic framework. Furthermore, this study contributes to improved genetic models for such giant ore deposits as well as more efficient exploration methodologies. In addition, better documentation of the Tambogrande deposits and tectonomagmatic history of the Lancones basin fills a significant void in the geological knowledgebase of one of the world’s major metallogenic belts, as well as one of the more poorly understood parts of the Andes. Chapter 1 Page 3 1.3 History Iron oxide occurrences were documented at Tambogrande nearly 100 years ago (Boletin de la Sociedad Geologica del Perú, 1904, in Tegart et al., 2000). But it was not until the mid‐
1970’s that interest was spurred again in the area by the joint venture of French state‐owned Bureau de Recherches, Geologiques et Minieres (BRGM) and the Instituto Geologico Minero y Metalurgico (INGEMMET) which found that the iron oxide (gossan) yielded anomalous base metal and silver values (Injoque et al., 1979). Susequent drilling of a self‐potential geophysical target in 1978 led to the discovery of massive sulphide mineralization at the TG1 deposit. Based on 21 diamond drill holes, the BRGM reported an inferred resource of 42.3 million tonnes grading 2.04% Cu, 1.45% Zn, 0.35% Pb and 38.4 g/t Ag (BRGM, 1981, in Tegart et al., 2000). Manhattan Minerals Corp. of Vancouver, Canada, became involved in the project in May, 1999. Re‐assaying of selected drill core suggested potential for the overlying oxide zone at TG1 to contain significant precious metals. A gravity survey, followed by additional drilling, resulted in an increased resource at TG1 as well as the discovery of two other massive sulphide deposits, TG3 and B5 (Table 2), approximately 0.5 and 12 km south of TG1, respectively (Tegart et al., 2000). An updated resource was defined for the TG1 sulphide deposit, and new resources for the TG1 oxide and TG3 sulphide deposits, were reported in 2002 (Table 1). However, due to socio‐political issues at the local and regional level, which ultimately led to the destruction of the company’s camp in early 2001, Manhattan lost title to the project in 2004. The Tambogrande deposits remain undeveloped and there is presently no known active exploration in the area. Chapter 1 Page 4 Detailed reports related to VMS deposits in northwestern Perú provide excellent documentation of the massive sulphide deposits but are limited in scope and have not evaluated the broader regional geologic setting (Tegart et al. 2000; Injoque et al., 1979). 1.4 VMS Deposit Classification and Genetic Models Volcanogenic massive sulphide deposits (Franklin et al., 1981; Lydon, 1984, 1988; Ohmoto, 1996; Barrie and Hannington, 1999), also commonly termed ‘volcanic‐hosted massive sulphide’ deposits (VHMS, Large, 1992), are ‘strata‐bound accumulations of sulphide minerals that precipitated at or near the sea floor in spatial, temporal, and genetic association with contemporaneous volcanism’ (Franklin et al., 2005). The broadly accepted genetic model, based mostly on empirical observations from ancient land‐based and modern seafloor systems, considers VMS deposits to be the products of district‐scale hydrothermal convection resulting from anomalous thermal input into the shallow crust (Fig. 1.2; Franklin et al., 2005). Down‐drawn seawater is the primary component of the hydrothermal fluid and heat‐
induced fluid‐rock reactions facilitate the transfer of metals and sulphur from the substrate into the convecting hydrothermal system (Franklin et al., 2005). Reduced seawater sulphate may constitute a minor sulphur component in VMS sulphides. Chloride complexing is considered responsible for most metal transport with precipitation triggered by cooling. To a much lesser amount bi‐sulphide complexes are thought to also play a role as transporting agent, especially for Au, with oxidation causing metal precipitation (Hannington et al., 1995). Mixing of a metal‐
bearing hydrothermal fluid with cool (± oxidized) seawater at or near the seafloor is the primary method for sulphide deposition (Franklin et al., 2005). Some VMS models invoke high‐level intrusive phases as heat sources and include magmatic‐derived fluid as an additional metals source (Galley, 1993; Franklin et al., 1981). VMS deposits are broadly syngenetic with the host Chapter 1 Page 5 rock assemblages, which vary from volcanic‐ to sediment‐dominated (Barrie and Hannington, 1999). Extensional basins are integral to nearly all VMS settings for two main reasons: (i) the heat transfer into the upper crust due to the intrusion and advection of magma into attenuated and extended crust; and (ii) the development of normal faults which channel hydrothermal fluids and focus the deposition of sulphide minerals. The geometries of VMS deposits vary significantly, but many occur as mounds on the seafloor, such as at the TAG hydrothermal mound (Fig. 1.3)(Herzig and Hannington, 1995). In the rock record, deposits may be bulbous to tabular lenses of massive sulphide, i.e., 60‐100% sulphide minerals (Lydon 1984), underlain by discordant zones of veined (‘stringer’ or ‘stockwork’) and disseminated mineralization. The discordant sulphide zones represent the paleo channel‐ways, usually synvolcanic faults, which facilitated metalliferous fluid transfer to the seafloor environment to form the VMS deposit. VMS deposits are known to occur from the Archean to modern times with notable spikes in the number of deposits in the Late Archean and Early Proterozoic (Fig. 1.4). VMS deposits are more evenly distributed throughout the Phanerozoic when compared to the Precambrian with the notable abundance of deposits in the early Paleozoic (Fig. 1.4). The Tambogrande deposits occur during a period of increased, but not prolific, VMS formation. Classification of massive sulphide deposits is most effective using the lithostratigraphy of the host assemblage (Barrie and Hannington, 1999; Franklin et al., 2005). Bimodal‐mafic deposits, which exemplify those at Tambogrande, are characterized by volcanic‐dominated stratigraphic sequences consisting largely of basalt and >3‐25% felsic volcanic rocks. Other examples of large VMS deposits from this class occur in the Archean, including the Noranda camp (Gibson and Watkinson, 1990) and Kidd Creek (Hannington and Barrie, 1999), the Proterozoic, such as Flin Chapter 1 Page 6 Flon (Syme and Bailes, 1993) and the Phanerozoic, such as the Urals deposits of Sibai and Gai (Herrington et al., 2005). In Latin America, significantly large bimodal‐mafic‐type Jurassic‐
Cretaceous VMS deposits occur at Cerro Lindo in central‐western Perú (Ly Zevallos, 2000) and at San Nicolas, central Mexico (Johnson et al., 2000; Danielson, 2000). Giant VMS Deposits The Tambogrande VMS district hosts three known deposits located near and to the south of the village of Tambogrande in the Piura district of northwestern Perú (Fig. 1.1). The Tambogrande deposits, with a collective tonnage of approximately 300 Mt, are within the upper 3% of all bimodal‐mafic type VMS deposits in terms of size and contained metal (Fig. 1.5; database from Franklin et al., 2005). The term ‘giant’ is applied to deposits with greater than 50 Mt of massive sulphide ore, whereas supergiants are often ascribed to deposits containing more than >100 Mt (Barrie and Hannington, 1999). The Tambogrande deposits are the largest of the Mesozoic massive sulphide deposits in Perú and represent the most significant group of VMS deposits of continental South America. Grade and tonnage estimates for the deposits are listed in Table 1. No resource estimate is available for B5, but the deposit had drillhole intersections of up to 280 m of massive sulphide at similar grades to TG1 and TG3 (Tegart et al., 2000). The TG1 deposit also has an associated supergene oxide zone (Table 1). 1.5 Controls on ‘Giant’ VMS systems Giant deposits vary in morphology, host rocks, mineralogy and metal content and such size may require either special processes or an ideal combination of processes in their formation. Two main factors must be considered for the formation of VMS deposits: (A) the nature/source of the hydrothermal fluids and (B) the depositional seafloor environment. With Chapter 1 Page 7 respect to the hydrothermal fluids, several variables could account for the delivery of a significant amount of metal to the seafloor environment. The possibilities include (i) a long‐
lived hydrothermal system allowing for the protracted delivery of metalliferous hydrothermal fluid to the site of deposition. For instance, based on heat and fluid‐flow modeling, Barrie et al. (1999) proposed that the supergiant Kidd Creek ore body (~200 Mt) could have formed in 650,000 years, whereas modern ocean ridge massive sulphide deposits (~1 Mt) are estimated to have formed in less than 100,000 years (Rona, 1988). (ii) Discharge rates at hydrothermal vents are also important and depend on a number of variables, namely the rate of heat loss (heat flux) and the amount of fluid available. Heat and fluid fluxes are largely dependent on the distance to the heat source (in a convection model) or the size of the intrusion (magmatic‐
hydrothermal and convection models). Ultimately, permeability is the main control on the fluid flux and this is likely to be controlled by both the style and degree of structural deformation as well as primary lithological features. (iii) Finally, metal concentrations in the hydrothermal fluids may be of critical importance. In a magmatic model, the metal concentration depends on the composition of the intrusion, metal partitioning into a volatile phase, and source of ligands to transport metals. In a convection system fluid‐rock reactions are of key importance, wherein the amount of rock available for leaching (size of the system), metal contents of the source rocks, and the effectiveness the fluids to scavenge metals must be considered. The depositional site for sulphides at or below the seafloor must also be conducive to efficient sulphide precipitation and preservation. Only a relatively small portion (<1–5%) of metals emitted from modern seafloor hydrothermal vents are actually precipitated (Converse et al., 1984; Feely et al., 1994). The remainder are vented into the ocean column and dispersed. Sub‐seafloor sulphide precipitation as replacement of relatively permeable strata is Chapter 1 Page 8 a common feature to VMS mineralization (Doyle and Allen, 2003) and has been recognized at the Horne (Kerr and Gibson, 1993) and Kidd Creek deposits (Hannington et al., 1999). Ambient oceanic environmental conditions are especially important in some VMS settings. Eastoe and Gustin (1996) suggest an association of Phanerozoic VMS deposits with black shale, and hence oceanic anoxia, due to the efficiency of sulphide precipitation and preservation in anoxic seawater environments. Saez et al. (1999) argue that many giant to supergiant VMS deposits in the Iberian Pyrite Belt (Portugal, Spain) are associated with black shale horizons formed in a euxinic environment and hence bacterial reduction occurred. Similarly, the supergiant Brunswick No. 12 deposit is considered to have been deposited within the deeper segments of a continental back‐arc rift basin where bottom waters were anoxic and reducing conditions ideal for VMS formation (Goodfellow and Peter, 1996). In the absence of anoxia, models involving siliceous ‘caps’ as protective barriers to oxidative waters have been proposed, such as Aljustrel deposit of the IPB (Barriga and Fyfe, 1988), as well as some modern seafloor deposits where silica‐producing bacteria have also been observed (Humphris et al., 1995). 1.6 Thesis Objectives Giant massive sulphide deposits at Tambogrande are an enigma in that few VMS districts contain such large deposits and rarely is there more than one giant deposit within a single camp. The Tambogrande VMS deposits and Lancones basin of northwestern Perú represent an opportunity to evaluate the conditions which permit the generation of giant VMS deposits. A main goal of this research is to place VMS deposits at Tambogrande within the regional geological context and to explain the accompanying tectonomagmatic scenario responsible for the development of the Lancones basin. Key to achieving these goals are (i) the development of a stratigraphic model for the volcanic sequences and (ii) an evaluation of the Chapter 1 Page 9 igneous petrological associations with the massive sulphide deposits. The latter will assist in understanding possible igneous controls on ore formation in VMS systems. In addition, due to the pristine preservation state of the VMS deposits and the excellent three dimensional geological database developed from diamond drill core logging, the thesis also aims to develop a better understanding of the local paleomorphological, volcanic and structural controls. Documenting the local and regional geological attributes of the Tambogrande VMS setting help establish the depositional environments and tectonic controls for giant massive sulphide deposits of this type, and establish a geologic framework for comparisons with similar VMS settings. 1.7 Methodology 1.7.1 Core Logging A total of 441 diamond drill holes and more than 80,000 m of core, mostly from Manhattan Minerals work at TG1, TG3, and B5, and several targets outside of the known deposits, were available for study during 2000. In 2002, additional drill core was also made available by BHP Billiton and Compañía de Minas Buenaventura from drill projects north of Tambogrande. The author had the opportunity to log and sample approximately half of the drill core from Manhattan’s work on the deposits and all available core from the region. 1.7.2 Regional Mapping Mapping at a scale of 1:100,000 was conducted during February to April, 2002, mostly north of Tambogrande (Fig. 1.1). Outcrop is generally poor in the vicinity of the deposits but improves to the north. Access to the east was limited by poor road conditions in rugged mountainous terrain, and was discouraged in several regions within the field area due to unstable political conditions. The area had previously been mapped at 1:100,000 by Reyes and Caldas (1987) as part of a large regional project. Chapter 1 Page 10 1.7.3 Geochronology U‐Pb age determinations of volcanic and intrusive rocks from the study area provide the first known radiometric ages for the region. Approximately 40 samples of felsic volcanic rocks from drill core and outcrops were processed with about 25% of the samples yielding sufficient zircons for U‐Pb isotope analysis. Samples were analyzed at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia, Vancouver, Canada, using Thermal Ionization Mass Spectrometry (TIMS). Additional samples were analyzed at Stanford University using Sensitive High Resolution Ion MicroProbe ‐ Reverse Geometry (SHRIMP‐RG). Several intrusive phases were also dated by step‐heated Ar‐Ar using the noble gas mass spectrometer (NGMS) technique at PCIGR. Details of the methodology, precision and accuracy are available in Appendices A & B and at http://www.eos.ubc.ca/research/pcigr/Instrumentation.htm and http://shrimprg.stanford.edu/. 1.7.4 Lithogeochemistry Samples were collected from drill core and outcrop for whole rock major, trace and rare earth element (REE) analysis. Limited lithogeochemical studies were reported from the region (Tegart et al., 2000). Analyses were carried out at ALS Chemex Laboratories, Vancouver, Canada and the Department of Earth Sciences at Memorial University, St. John’s, Canada, using a combination of inductively coupled plasma – mass spectrometry (ICP‐MS) and inductively coupled plasma ‐ atomic emission spectrometry (ICP‐AES). Details of the methodology, precision and accuracy are available in Appendix C and at http://www.alschemex.com/learnmore/learnmore‐techinfo‐multielement‐wholerock.htm and http://www.mun.ca/earthsciences/ICPMS/Solution_ICP‐MS.php. Chapter 1 Page 11 1.7.5 Isotope Chemistry Pb, Sm‐Nd and Rb‐Sr isotope geochemical data were utilized as tracers of relative mantle and continental crust contributions to the volcanic rocks. Pb isotope analyses of ore samples were also carried out for VMS deposits in order to ascertain the source of the Pb (and by inference other metals). All analyses were carried out at the PCIGR. Details of the methodology are available at http://www.eos.ubc.ca/research/pcigr/Instrumentation.htm. 1.8 Presentation This study is presented as a series of research manuscripts for the purpose of publication in refereed professional journals pertaining to economic geology. As such, some repetition of material is inevitable. Chapter 2 ‐ Volcanic stratigraphy and geochronology of the Cretaceous Lancones basin, northwestern Perú – is a review of the regional geology of the Lancones basin with emphasis of the volcanic successions and is based on geology interpreted from mapping and diamond drill core logging. The chapter provides a report of U‐Pb zircon ages for felsic volcanic rocks throughout the volcanic arc sequence of the Lancones Basin, as well as U‐Pb zircon ages for some intrusive phases within the sequence. These are the first reported radiometric ages for the rocks of this region and will assist in (ii) developing the stratigraphic models and understanding the overall geological evolution of this region, such as the timing of volcanism and duration of the volcanic arc sequence, and (ii) constraining the timing of formation of VMS deposits. Pinpointing the precise time interval of ore formation is critical to establish temporal links with broader tectonic and magmatic events. Moreover, this chapter makes a significant contribution to the overall understanding of the tectonomagmatic framework and Mesozoic evolution of the northern Andes of Perú by documenting an area that has not been studied in detail. Chapter 1 Page 12 Chapter 3 – A reconstructed Cretaceous depositional setting for giant volcanogenic massive sulfide deposits at Tambogrande, northwestern Perú – illustrates the architecture of the depositional setting and VMS deposits and provides a detailed analysis of the immediate VMS‐hosting volcanic succession as well as reconstructed paleo‐sea‐floor models. This chapter utilizes geologic plans, sections and three‐dimensional perspectives to recreate the submarine structural and volcanic setting of the TG1 and TG3 deposits and emphasize the important role of the local submarine volcanic environment in the location and morphology and possibly the size of the VMS deposits. Tambogrande represents a rare occurrence of well preserved, giant VMS deposits and may serve as a physical model for similar deposits in deformed terrains. Chapter 4 ‐ Volcanic rock geochemistry and the tectonomagmatic setting of VMS deposits at Tambogrande, Perú ‐ this chapter evaluates the volcano‐stratigraphic lithochemical variations throughout the Lancones basin and demonstrates generally ‘oceanic’ arc‐like geochemical signatures that vary from tholeiitic to weakly calc‐alkaline. The ocean‐like affinities reaffirm the tectonic setting as a volcanic arc developed on very thin crust at the leading edge of continental South America. The data also invoke partial melting of the crust as the source for felsic volcanic rocks in the VMS environment, consistent with a high heat flow regime and favourable for VMS formation. Chapter 5 – Pb‐Sr‐Nd isotope systematics of Cretaceous arc volcanic rocks in the Lancones Basin near Tambogrande, Perú – Implications for VMS deposit models – this chapter is a tracer isotope study of the volcanic rocks hosting VMS deposits at Tambogrande and supports the conclusions of Chapter 4 regarding the tectonic setting and petrochemical evolution of the volcanic rocks in the Lancones basin. Specifically, the data support a mantle‐wedge source for basalts but indicate continental crust was a factor in the generation of the felsic volcanic rocks. Chapter 1 Page 13 The data also permit tests of models relating igneous petrochemistry to VMS deposits and suggests felsic volcanic rocks do not contribute to metals in the formation of the VMS deposits, but are related to the same thermal anomaly which enabled VMS formation. Chapter 6 – Summary, Discussion and Unresolved Questions ‐ A summary of the major findings of this study with some discussions and elaboration on the on key ideas. The key contributions of this research project include the following: (i) a comprehensive baseline documentation of an economically significant VMS district; (ii) a better understanding of the Cretaceous geology of the northwest Perú Andean segment and of VMS metallogenesis within the marginal basins of Perú; (iii) the reconstruction of the paleo‐depositional environment for massive sulphide mineralization at Tambogrande provides a unique perspective as many VMS deposits are more deformed and more difficult to reconstruct; (4) a synopsis of a continental‐
margin VMS setting with strong constraints on the relationship to global plate tectonics; and (5) more constrained models for VMS genesis in terms of the relationship to associated felsic volcanic rocks, metal sources and hydrothermal systems. The chapter also proposed topics for future work. Chapter 1 Page 14 Figure 1.1. Location maps and simplified geology for the study area. The locations of VMS deposits (TG1, TG3, and B5) in the Tambogrande area are also shown and field area of this study outlined. Geology modified after Jaillard et al. (1999) and Tegart et al. (2000). Chapter 1 Page 15 Figure 1.2. Schematic model for the formation of VMS deposits (from Franklin et al., 2005). Chapter 1 Page 16 Figure 1.3. Schematic section and model of a typical volcanogenic massive sulphide deposit from modern mid‐ocean ridge settings; after Herzig and Hannington (1995). Chapter 1 Page 17 Figure 1.4. Histogram of ages for global bimodal‐mafic type VMS deposits (n=327); data from Franklin et al. (2005). Chapter 1 Page 18 Figure 1.5. Metals versus size of the deposit (tonnes) for global VMS deposits of the bimodal‐
mafic class (n=326; data from Franklin et al., 2005). Tambogrande deposits are labeled. KC = Kidd Creek deposit. A. Copper and B. Zinc Chapter 1 Page 19 Figure 1.6. Gold grade (grams/tonne) versus size of the deposit (tonnes) for global VMS deposits of the bimodal‐mafic class (n=326; data from Franklin et al., 2005). Tambogrande deposits are labeled. KC = Kidd Creek deposit. Chapter 1 Page 20 Table 1.1. Individual deposit data from number drill holes, tonnage and grade (Manhattan Minerals, 2002). Deposit
# Drill Category1
Metric
Cu % Zn % Au g/t Ag g/t
holes
Tonnes
Indicated2
56,156,000
1.6
1
0.5
26
2
Inferred
3,295,000
1.5
0.8
0.4
18
TG1
3
357
sulphide
Probable
49,200,000 1.6
1
0.4
18
ore
All
108,651,000 1.6
1.0
0.5
22
categories
357
Indicated4
Inferred4
Probable5
All
categories
7,964,000
725,000
8,056,000
16,745,000
-------
-------
3.6
3.4
3.5
3.5
62
62
67
64
TG3
sulphide
ore
53
Inferred6
82,000,000
1
1.4
0.8
25
B5
sulphide
ore
14
?
?
?
?
?
TG1 oxide
ore
1 The reserves and resources are calculated in accordance with National Instrument 43‐101 ‐ Standards of Disclosure For Mineral Projects. Mineral Reserves presented here are as at December 31, 2000, based on pre‐
feasibility Studies. Mineral Resources for TG‐1 have been reviewed and updated (March 2002). All calculations are based on the following metal prices: Cu US$ 0.90/lb, Zn US$ 0.55/lb, Au US$ 300/lb, Ag US$ 5.00/lb. 2 resource with 0.75% Cu equivalent cutoff (Cu equiv. = Cu% + 0.61 Zn%) for sulphide ore and 1 g/t Au cutoff for oxide ore. 3
reserves are calculated on variable NSR cut‐offs and incorporate 200,000 tonnes of external dilution and 400,000 tonnes of internal dike dilution, both taken at 0 grade. 4 resource is calculated based on a cut‐off grade of 1.0 g/t gold. High values have been cut to 20 g/t Au and 150 g/t Ag in the lower grade Oxide zone (main body) and 50 g/t Au and 2000 g/t Ag in the higher‐grade Transition zone (contact zone with sulphide). 5
reserves are calculated on a Net Smelter Return (NSR) cut‐off of US$8.53 and incorporate 95,000 tonnes of external dilution taken at zero grade. 6 resource with cutoff at 0.5% Cu equivalent. Note: lead assays were not reported in resource estimates but for approximation purposes yield ~0.11% Pb at TG1, based on the author’s calculations in this study for more than 8600 ore grade samples (i.e., 0.75% Cu equivalent). Chapter 1 Page 21 1.9 References Atherton, M.P., Pitcher, W.S., and Warden, V. 1983. The Mesozoic marginal basin of central Peru, Nature, 305: 303‐306. Barrie, C.T. and Hannington, M.D. 1999. Classification of volcanic‐associated massive sulphide deposits based on host‐rock composition. In Volcanic‐associated massive sulphide deposits: processes and examples in modern and ancient settings. Edited by C.T. Barrie and M.D. Hannington. Reviews in Economic Geology, 8: 1‐11. Barrie, C.T., Cathles, L.M., Erendi, A. 1999. Finite element heat and fluid‐flow computer simulations of a deep ultramafic sill model for the giant Kidd Creek volcanic‐associated massive sulphide deposit, Abitibi Subprovince, Canada. In The Giant Kidd Creek Volcanogenic Massive Sulphide Deposit, Western Abitibi Subprovince. Edited by M.D. Hannington and C.T. Barrie. Economic Geology Monograph 10, pp. 529‐540. Barriga, F.J.A.S. and Fyfe, W.S. 1988. Giant pyritic base‐metal deposits: the example of Feitais (Aljustrel, Portugal). Chemical Geology, 69: 331‐343. Benavides‐Caceres, V. 1999. Orogenic evolution of the Peruvian Andes; the Andean Cycle; geology and ore deposits of the central Andes, Special Publication ‐ Society of Economic Geologists, 7: 61‐107. Clark, A.H. 1993. Are outsize porphyry copper deposits either anatomically or environmentally distinctive?; giant ore deposits. In Giant Deposits. Edited by Whiting, B.H., Hodgson, C.J., and Mason, R. Society of Economic Geologists, Special Publication 2, pp. 213‐284. Clark, A.H. 1995. Proceeding of the Second Giant Ore Deposits Workshop, Giant Ore Deposits‐
II; Controls on the scale of orogenic magmatic‐hydrothermal mineralization: Department of Geological Sciences, Queen's University, Kingston, Ontario, Canada, 753 p. Converse, D.R., Holland, H.D., and Edmond, J.M. 1984. Flow rates in the axial hot springs of the East Pacific Rise (21oN): implications for the heat budget and the formation of massive sulphide deposits. Earth and Planetary Science Letters, 69: 159‐175. Dalziel, I.W.D., Vine, F.J. and Smith, A.G. 1981. Back‐arc extension in the southern Andes; a review and critical reappraisal; extensional tectonics associated with convergent plate boundaries, Philosophical Transactions of the Royal Society of London, Series A: Mathematical and Physical Sciences, 300: 319‐335. Danielson, T., J. 2000. Age, paleotectonic setting, and common Pb isotope signature of the San Nicolás volcanogenic massive sulphide deposit, southeastern Zacatecas state, central Mexico. Unpublished M.Sc. thesis, University of British Columbia, Vancouver, B.C., Canada, 120 p. Chapter 1 Page 22 Doyle, M. G., and Allen, R. L. 2003. Subseafloor replacement in volcanic‐hosted massive sulfide deposits. Ore Geology Reviews, 23: 183‐222. Eastoe, C.J. and Gustin, M.M. 1996. Volcanogenic massive sulfide deposits and anoxia in the Phanerozoic oceans, Ore Geology Reviews, 10: 179‐197. Feely, R.A., Massoth, G.J., Trefry, J.H., Baker, E.T., Paulson, A.J., and Lebon, G.T. 1994. Composition and sedimentation of hydrothermal plume particles from the North Cleft segment, Juan de Fuca Ridge. Journal of Geophysical Research, 99: 4985‐5006. Franklin, J.M., Sangster, D.M., and Lydon, J.W. 1981. Volcanic‐associated massive sulfide deposits. In Economic geology; Seventy‐fifth Anniversary Volume; 1905‐1980. Edited by B.J. Skinner. Economic Geology Publishing Co., pp. 485‐627. Franklin, J.M., Gibson, H.L., Jonasson, I.R., and Galley, A.G. 2005. Volcanogenic massive sulfide deposits. In Economic Geology; one hundredth anniversary volume, 1905‐2005. Edited by J.W. Hedenquist, J.F.H. Thompson, R.J. Goldfarb and J.P. Richards. Society of Economic Geologists, pp. 523‐560. Galley, A.G. 1993. Semi‐conformable alteration zones in volcanogenic massive sulphide districts. Journal of Geochemical Exploration, 48: 175‐200. Gibson, H.L. and Watkinson, D.H. 1990. Volcanogenic massive sulphide deposits of the Noranda cauldron and shield volcano, Quebec; Special Volume ‐ Canadian Institute of Mining and Metallurgy, 43: 119‐132. Gibson, H.L., Kerr, D.J., and Cattalani, S. 2000. The Horne Mine; geology, history, influence on genetic models, and a comparison to the Kidd Creek Mine, Exploration and Mining Geology, 9: 91‐111. Goodfellow, W.D. and Zierenberg, R.A. 1999. Genesis of massive sulfide deposits at sediment‐
covered spreading centers, Reviews in Economic Geology, 8: 297‐324. Goodfellow, W.D. and Peter, J.M. 1996. Sulphur isotope composition of the Brunswick No. 12 massive sulphide deposit, Bathurst Mining Camp, New Brunswick: implications for ambient environment, sulphur source, and ore genesis. Canadian Journal of Earth Sciences, 33: 231‐251. Hannington, M.D. and Barrie, C.T. 1999. The Giant Kidd Creek Volcanogenic Massive Sulphide Deposit, Western Abitibi Subprovince, Canada. Economic Geology, Monograph 10, 672 p. Hannington, M.D., Jonasson, I.R., Herzig, P.M., and Peterson, S. 1995. Physical and chemical processes of seafloor mineralization at mid‐ocean ridges. American Geophysical Union Monograph 91, pp. 115‐157. Chapter 1 Page 23 Hannington, M.D., Bleeker, W., Kjarsgaard, I. 1999. Sulfide mineralogy, geochemistry and ore genesis of the Kidd Creek deposit: part I. North, central, and south orebodies. In The Giant Kidd Creek Volcanogenic Massive Sulphide Deposit, Western Abitibi Subprovince. Edited by M.D. Hannington and C.T. Barrie. Economic Geology Monograph 10, pp.163‐
224. Hanson, R.E., Wilson, T.J., Harmon, R.S. and Rapela, C.W. 1991. Submarine rhyolitic volcanism in a Jurassic proto‐marginal basin; southern Andes, Chile and Argentina; Andean magmatism and its tectonic setting, Special Paper ‐ Geological Society of America, 265: 13‐27. Hedenquist, J.W., Arribas R, A., and Gonzalez‐Urien, E. 2000. Exploration for epithermal gold deposits; Gold in 2000, Reviews in Economic Geology, 13: 245‐277. Herrington, R., Maslennikov, V., Zaykov, V., Seravkin, I., Kosarev, A., Buschmann, B., Orgeval, J., Holland, N., Tesalina, S., Nimis, P., and Armstrong, R. 2005. Classification of VMS deposits; lessons from the Uralides; Ore Geology Reviews, 27: 203‐237. Herzig, P.M. and Hannington, M.D. 1995. Polymetallic massive sulfides at the modern seafloor, a review; Ore Geology Reviews, 10: 95‐115. Humphris, S.E., Zierenberg, R.A., Mullineaux, L.S., and Thomson, R.E. 1995. Seafloor Hydrothermal Systems: Physical, Chemical, Biological, and Geological Interactions. Geophysical Monograph 91, American Geophysical Union, Washington, D.C., 466 p. Injoque, J. Miranda, C., Duninn‐Borkowski, E. 1979. Estudio de la genesis del yacimiento de Tambo Grande y sus implicancias. Boletin de la Sociedad del Peru, 64: 73‐99. Jaillard, É, Laubacher, G., Bengston, P., Dhondt, A., and Bulot, L. 1999. Stratigraphy and evolution of the Cretaceous forearc “Celica‐Lancones basin” of Southwestern Ecuador. Journal of South American Earth Sciences, 12: 51‐68. Jaillard, E., Herail, G., Monfret, T., Diaz‐Martinez, E., Baby, P., Lavenu, A. and Dumont, J.F. 2000. Tectonic evolution of the Andes of Ecuador, Perú, Bolivia and northernmost Chile; tectonic evolution of South America. In Tectonic evolution of South America, 31st International Geological Congress, Rio de Janeiro, Brazil. Edited by U.G. Cordani, E.J. Milani, A. Thomaz Filho and D.A. Campos, pp. 481–559. Jaillard, E., Ordonez, M., Berrones, G., Bengtson, P., Bonhomme, M., Jimenez, N., and Zambrano, I. 1996. Sedimentary and tectonic evolution of the arc zone of southwestern Ecuador during Late Cretaceous and Early Tertiary times; Andean geodynamics. Journal of South American Earth Sciences, 9: 131‐140. Johnson, B.J., Montante‐Martinez, J.A., Cenela‐Barboza, M., and Danielson, T.J. 2000. Geology of the San Nicolás deposit, Zacatecas. In VMS Deposits of Latin America. Edited by R. Sherlock and M.A.V. Logan. Geological Association of Canada, Mineral Deposits Division, Special Paper No.2., pp. 71‐85. Chapter 1 Page 24 Kerr, D. and Gibson, H.L. 1993. A comparison of the Horne volcanogenic massive sulphide deposit and intracauldron deposits of the Mine sequence, Noranda, Quebec. Economic Geology, 88: 1419‐1442. Large, R.R. 1992. Australian volcanic‐hosted massive sulfide deposits: Features, styles, and genetic models. Economic Geology, 87: 471–510. Lydon, J.W. 1984. Volcanogenic massive sulphide deposits, part 1: A descriptive model. Geoscience Canada, 15: 195‐202. Lydon, J.W. 1988. Volcanogenic massive sulphide deposits, part 2: genetic models. Geoscience Canada, 15: 43‐65. Ly Zevallos, P. 2000. Cerro Lindo Project. In VMS Deposits of Latin America. Edited by R. Sherlock and M.A.V. Logan. Geological Association of Canada, Mineral Deposits Division, Special Paper No.2., pp. 407‐422. Manhattan Minerals Corporation. 2002. Annual Report Mitouard, P., Kissel, C., and Laj, C. 1990. Post‐Oligocene rotations in southern Ecuador and northern Peru and the formation of the Huancabamba deflection in the Andean cordillera, Earth and Planetary Science Letters, 98: 329‐339. Mourier, T., Laj, C., Mégard, F., Roperch, P., Mitouard, P., and Farfan Medrano, A. 1988. An accreted continental terrane in northwestern Peru, Earth and Planetary Science Letters, 88: 182‐192. Ohmoto, H. 1996. Formation of volcanogenic massive sulfide deposits; the Kuroko perspective; Ore Geology Reviews, 10: 135‐177. Reyes, L.R. and Caldas, J.Y. 1987. Geologia de los Cuadranglos de las Playas, La Tina, Las Lomas, Ayabaca, San Antonio. Instituto Geologico Minero y Metalurgio, Bul. 49., 83 p. Rona, P. 1988. Hydrothermal mineralization at ocean ridges; Canadian Mineralogist, 26: 431‐
465. Sáez, R., Pascual, E., Toscano, M. and Almodóvar, G.R. 1999. The Iberian type of volcano‐
sedimentary massive sulphide deposits. Mineralium Deposita, 34: 549‐570. Sangster, D.F. 1972. Precambrian massive sulphide deposits in Canada: a review. Geological Survey of Canada, Paper 72‐22. Singer, D.A. 1995. World class base and precious metal deposits; a quantitative analysis, Economic Geology and the Bulletin of the Society of Economic Geologists, 90: 88‐104. Chapter 1 Page 25 Syme, E.C., and Bailes, A.H., 1993. Stratigraphic and tectonic setting of Early Proterozoic volcanogenic massive sulphide deposits, Flin Flon, Manitoba: Economic Geology, 88: 566‐
589. Tegart, P., Allen, G., Carstensen, A. 2000. Regional setting, stratigraphy, alteration and mineralization of the Tambo Grande VMS district, Piura Department, Northern Perú. In VMS Deposits of Latin America. Edited by R. Sherlock and M.A.V. Logan. Geological Association of Canada, Mineral Deposits Division, Special Paper No.2. pp. 375‐405. Vergara, M., Levi, B., Nyström, J.O., Cancino, A. 1995. Jurassic and Early Cretaceous island arc volcanism, extension, and subsidence in the Coast Range of central Chile. Geological Society of America Bulletin, 107: 1427‐1440. Chapter 1 Page 26 Chapter 2. Volcanic Stratigraphy and Geochronology of the Cretaceous Lancones Basin, Northwestern Perú 1 2.1 Overview A ~10 km‐thick sequence of Cretaceous basaltic to rhyolitic volcanic rocks forms the arc component of the Lancones Basin in northwestern Perú and underlies part of the Huancabamba deflection. The marine volcanic successions show markedly different compositional features and depositional facies and suggest two main formational environments, consistent with a maturing arc and shallowing marine basin. The earliest volcanism accompanying rifting, referred to as phase 1, was dominated by basaltic pillow lava and breccias with lesser aphyric to feldspar‐quartz porphyritic felsic volcanic rocks. These volcanic successions filled the lowest exposed portion of the basin and were accompanied by volcanogenic massive sulphide (VMS) deposits, which are inferred to have formed in a relatively deep marine setting. U‐Pb zircon dating of felsic volcanic rocks associated with VMS deposits at Tambogrande indicate ages from 104.8 ±1.3 to 100.2 ±0.5 Ma for the phase 1 volcanic sequence. The timing of onset of rift‐related volcanism is not well constrained, but is therefore of middle Albian age or older. Phase 2 volcanism is composed of latest Albian to Turonian successions of relatively more felsic‐rich volcaniclastic rocks and yields U‐Pb zircon ages of 99.3±0.3 to 91.1 ±1.0 Ma. Successions of phase 2 volcanism are intercalated and overlain by siliciclastic and carbonate sedimentary sequences prevalent in the western forearc section of the basin. Phase 2 volcanism is followed by granitoid plutonism of the Coastal Batholith beginning in the Late Cretaceous. 1
A version of this chapter will be submitted for publication. Winter, L.S., Tosdal, R., and Mortensen, J. Volcanic
Stratigraphy and Geochronology of the Cretaceous Lancones Basin, Northwestern Perú.
Chapter 2 Page 27 The genesis of the Cretaceous Lancones Basin and other equivalent rift‐related basins in western South America, including the Western Peruvian Trough, is related tectonically to the break‐up of Gondwana. Phase 1 volcanism in the Lancones Basin in Albian times coincided with the initial rift stage, prior to active oceanic spreading, between South America and Africa. During this time the relatively stationary western margin of continental South America was undergoing extension and rifting due to a westwards (oceanwards) retreating arc, resembling a Mariana arc type setting. The Mochica orogeny marks the termination of rifting, subsidence and related volcanism along the western margin of South America. This orogenic event also broadly coincides with the onset of spreading of the South Atlantic and westward drift of the South American continent. Subsequent phase 2 volcanism was more continental arc‐like under an Andean‐type arc scenario. 2.2 Introduction The Lancones Basin of northwestern Perú (also known as the Celica‐Lancones Basin; Jaillard et al., 1999), represents a Cretaceous marine volcanic and sedimentary succession with large tonnage deposits of base‐ and precious‐metal‐bearing massive sulphides around Tambogrande (Tegart et al., 2000; Chapter 3). The basin overlaps both the Western Cordillera and the Para‐
Andean Depression of northwestern Perú and southwestern Ecuador (Fig. 2.1) and represents the northernmost of a series of Late Jurassic to Early Cretaceous continental margin, arc‐related rift basins extending along the South American western margin through Perú (Huarmey and Cañete Basins or Western Peruvian Trough; Myers, 1974; Cobbing et al., 1981; Atherton et al., 1983), Chile (Coast Range; Vergara et al., 1995), and Argentina (Rocas Verdes; Dalziel, 1981; Hanson and Wilson, 1991). In comparison to the other segments of the marginal rift system, Chapter 2 Page 28 the Lancones basin is well preserved and includes the arc and forearc components as well as accreted allochthonous crustal blocks (i.e., the Amotape terrain; Fig. 2.2). In this chapter the geology of the Lancones Basin, with emphasis on the volcanic successions and the eastern portion of the basin, are reviewed based on recent mapping and examination of diamond drill core. This documentation represents a detailed account of the volcanic successions of the Lancones basin and places the Tambogrande volcanogenic massive sulphide (VMS) deposits in a stratigraphic context. These VMS deposits represent the largest of the massive sulphide deposits in Perú and represent the most significant group of VMS deposits in South America, each with a gross tonnage of massive sulphide of ~100 Mt. Furthermore, they are within the upper 3% of all VMS deposits of their type globally with respect to size (Franklin et al., 2005). In addition, the first U‐Pb zircon ages for the volcanic rocks from the Lancones marginal basin of northwestern coastal Perú are presented herein. In this paper (i) a chronostratigraphic framework for the volcanic successions of the Lancones basin is provided; (ii) the field criteria are used to recognize the various sequences and redefine formations; (iii) the volcanic depositional setting including the VMS environs and the evolution of the setting through time are presented; (iv) a comparison of the volcanic stratigraphy of the Lancones basin with other Cretaceous sequences in Perú and elsewhere is provided; and (v) an explanation of how the Mesozoic volcanic arc is related to the tectonomagmatic evolution of South America is provided. Because northwestern Perú lies in a major oroclinal bend, the Huancabamba deflection, additional constraints on the tectonic development of this region improves upon understanding of the tectonomagmatic history the Andes. Chapter 2 Page 29 2.3 Tectonic Setting The tectonic evolution of the Jurassic to Tertiary South American western margin was largely a function of terrain accretion and variable plate convergence directions and rates (Soler and Bonhomme, 1990). The latter was influenced by the late stages of Gondwana breakup in the Early Cretaceous. These events which triggered subduction along the western margin of the continent mark the oldest phase of the Andean Cycle (Benavides‐Cáceres, 1999). Throughout the Jurassic, a southeast‐directed subduction system was responsible for continental arc volcanism along the Ecuadorian segment (Litherland et al., 1994), whereas a sinistral transform system dominated the Peruvian segment (Fig. 2.3A; Jaillard et al., 2000). A shift towards northeast‐directed convergence occurred in the Early Cretaceous. This is indicated by the termination of the arc along the Ecuadorian segment. The Amotape terrane is a microcontinental block of Paleozoic or older metasedimentary rocks and Triassic metaplutonic rocks based on U‐Pb zircon ages (Noble et al., 1997; Appendix B). Within the Amotape Terrane in southern Ecuador, high pressure metamorphosed oceanic rocks yield cooling ages of ~132‐110 Ma (Arculus et al., 1999; Bosch et al., 2002) that record accretion to continental South America during the Neocomian. The allochthonous Amotape terrane was transported northward and accreted in the Early Cretaceous with northeast‐
trending dextral faults developed during clockwise rotation (Mourier et al., 1988; Fig. 2.3B). The accretion is temporally linked to, and likely triggered, the westward relocation of the plate boundary which manifested as a new subduction zone along the north‐northwest‐trending Peruvian segment. Under this Mariana‐type arc system, steep subduction and slab roll‐back caused extension and attenuation in the overriding continental plate and resulted in rifting and the formation of the Lancones Basin, followed by the deposition of marine sequences and the Chapter 2 Page 30 eruption of large volumes of mafic‐dominated arc volcanic rocks (Fig. 2.3C; Benavides‐Cáceres, 1999). A clockwise rotation may have been related to the opening of the Lancones Basin in Albian times (Winter et al., 2002). Gravity modeling of crustal structure along the Peruvian continental margin indicates that a dense arch‐like structure of 3.0 g/cm3 underlies the volcanic‐dominated portion of the Western Peruvian Trough and possibly represents the intrusion of basic material into the continental crust (Jones, 1981). In late Albian times the geodynamical cycle shifted towards Andean‐type subduction and marked the first Andean compressive tectonism, i.e., Mochica Phase (Mégard, 1984) and subsequent continental arc volcanism and plutonism (Coastal Batholith). An increasing convergence rate through the Albian (Soler and Bonhomme, 1990), temporally linked to the opening of the South Atlantic, may have been responsible for this transition in subduction zone setting. A series of post‐rift collisional events along the Northern Andes in Ecuador shortened the Lancones basin and contributed an additional component of clockwise rotation (Mitouard et al., 1990). The accretion of the Pallatanga Terrane in the Late Cretaceous to Paleocene (Fig. 2.3D) and the Macuchi Island Arc in the late Eocene to early Oligocene (Huges and Pilatasig, 2002; Spikings et al., 2005) led to the current terrain configuration (Fig. 2.3E). This compressive tectonic regime continues to the present day. 2.4 Regional Geology The Lancones Basin is situated at a major oroclinal bend in the Andes, the Huancabamba deflection, which separates the north‐northwest‐trending Peruvian Andes from the northeast‐
trending Ecuadorian Andes (Mitouard et al., 1990; Fig. 2.2). The basin is limited to the east‐
southeast and southwest to north by continental crustal blocks that represent the Jurassic to Early Cretaceous pre‐rift Andean margin and were topographical highs during deposition in Chapter 2 Page 31 Mesozoic times (Cobbing et al., 1981). To the southeast, the Paleozoic(?) Olmos Massif is a probable reactivated margin of the Amazonian craton (Macfarlane, 1999). This poorly defined terrain consists of pre‐Ordovician greenschist facies pelitic to psammitic rocks overlain by platform carbonate rocks of Triassic to Early Jurassic age, considered equivalent to the Marañon Geanticline farther southeast in the Perú (Cobbing et al., 1981; Reyes and Caldas, 1987; Mourier et al., 1988; Litherland et al., 1994). Bordering the Lancones basin to the southwest, west and north, are Paleozoic or older meta‐sedimentary rocks and Triassic granitic rocks of the Amotape Range (Mourier et al., 1988; Aspden et al., 1995; Noble et al., 1997; Appendix B). Volcanic and sedimentary rocks of the Cretaceous Lancones basin can be subdivided into an eastern volcanic arc and western forearc. The rocks are exposed over 135 kilometres in strike and ~150 kilometres width through northwestern Perú and southwestern Ecuador. The basin extends beneath Tertiary cover in the southwest for up to an additional 50 kilometres (Fig. 2.2). An eastern volcanic arc sequence up to 80 km wide consists of mafic to felsic volcanic and volcaniclastic rocks. These uppermost of the volcanic successions grade into sedimentary rocks which dominate the western forearc portion of the Lancones Basin (Jaillard et al., 1999). The forearc turbiditic subbasin was filled with the 3 km thick Copa Sombrero Group that interfingers at the base of the group but, for the most part, overlap and buried the volcanic arc sequence (Chávez and Nuñez del Prado, 1991; Morris and Aleman, 1975; Jaillard et al., 1996, 1999). Late Cretaceous‐Tertiary marine sequences as well as Pleistocene and recent sediments unconformably cover all older rocks. 2.5 Volcanic Stratigraphy The stratigraphic units that define the volcanic arc sequence of the Lancones Basin include a wide spectrum of compositions and volcanic rock types ranging from effusive lava flows to Chapter 2 Page 32 pyroclastic rocks, from mafic to felsic volcanic rocks with less abundant intermediate compositions, and with variable proportions of intercalated sedimentary rocks. In general, the sequence evolved from lava flow facies to volcaniclastic‐rich to sedimentary successions. The sequence also appears to evolve from deep to shallow marine, and with possibly subaerial volcanic rocks deposited in the uppermost sections. Four main formations (modified from Reyes and Caldas, 1987) define the volcanic arc sequence of the Lancones Basin (Figs. 4, 5). The Cerro San Lorenzo Formation is introduced in this study as a new formation with the type locality located south of the San Lorenzo reservoir. This formation represents a significant component of, but is considered distinctive from, the former Cerro El Ereo Formation of Reyes and Caldas (1987) within which it was previously included. The Cerro El Ereo Formation has been retained, named for the topographically Ereo Hill which is comprised wholly of this sequence. Cerro El Ereo is defined by a distinctive porphyritic mafic lava flow sequence that differs from the pillow basalt sequence of the Cerro San Lorenzo Formation in terms of composition and depositional facies. In addition, the La Bocana and Lancones Formations have been revised to define the La Bocana Formation as a volcanic‐volcaniclastic sequence, whereas the overlying and thicker calcareous‐siliciclastic sedimentary dominated successions are assigned solely to the Lancones Formation. Furthermore, estimates from this study suggest a significantly greater total thickness (~ 8 to 10 km; Fig. 2.6) for the volcanic arc sequence as compared to previous estimates of ~2 km (Reyes and Caldas, 1987). In Ecuador the volcanic and volcaniclastic sequence has not been studied in detail and is described as a 2 to 3 km‐thick package of dominantly mafic pillow lavas and related volcaniclastic rocks (Jaillard et al., 1996). Fossils have been reported only from the La Bocana Formation and stratigraphically higher successions (Reyes and Caldas, 1987). Chapter 2 Page 33 The volcanic arc sequence of the Lancones Basin is subdivided into two main tectono‐
volcanic phases based on depositional facies, composition and chronology. The Cerro San Lorenzo Formation represents phase 1 and the Cerro El Ereo, La Bocana and Lancones Formations compose phase 2. Phase 1, a mafic‐dominated sequence characterized by lava flows and associated breccia, minor aphyric felsic lava, and general absence of siliciclastic sedimentary rocks is interpreted to represent a deep water environment. The basal contact has not been seen in outcrop or drill core but siliciclastic rocks of the San Pedro Group present to the east are interpreted as the earliest of the Cretaceous sequences within the basin (Reyes and Caldas, 1987). Geochronological data presented herein limit phase 1 volcanic rocks to the late Albian. The phase 1 volcanic sequence is of economic interest as it hosts all known VMS deposits in the Lancones basin. The phase 2 volcanic cycle is defined as an 8 km‐thick sequence of mafic to felsic volcanic and volcaniclastic rocks with subordinate sedimentary rocks that represent a more shallow water setting. The upper part of the phase 2 volcanic cycle, the Lancones Formation, is dominated by volcaniclastic rocks and intercalated with sedimentary and notably calcareous rocks that mark a transition to forearc turbidite of the Copa Sombrero Group. Rocks within phase 2 are late Albian to Turonian(?) based on geochronological data presented in this chapter. Metamorphic grades range from zeolite to lower greenschist facies, with the higher metamorphic grade mineral assemblages due to the thermal metamorphism around younger plutonic rocks. Diagenesis and seafloor metasomatism,due to seawater‐rock interaction, resulted in a wide range of low‐temperature replacement and open‐space‐filling minerals within basaltic rocks, including analcite, albite (after plagioclase), amphibole (uralitized Chapter 2 Page 34 clinopyroxene), carbonate minerals, chlorite, epidote, hematite, palagonite (after groundmass glass), prehnite, pumpellyite, sericite or various clays (sausseritized feldspar), and zeolites. Hydrothermal alteration is confined to discordant zones of the footwall rocks immediately below massive sulphide deposits and includes variable replacement of the rocks with Fe‐
chlorite, sericite and quartz in addition to stringer and disseminated sulphide mineralization (Tegart et al., 2000). None of the rocks in this study show the effects of significant ductile strain and primary textures are generally well preserved. 2.5.1 Cerro San Lorenzo Formation The lowermost of the volcanic sequences, the Cerro San Lorenzo Formation (Fig. 2.6), is characterized by bimodal volcanic rocks dominated by pillow basalt. The depositional environment is considered to have been relatively deep marine, based on the absence of the pyroclastic rocks and with sedimentary rocks limited to thin units of laminated black mudstone. Felsic volcanic rocks represent approximately 10% or less of the total volume of the formation. Drill holes in the south of the basin near the vicinity of the massive sulphide deposits tested to a maximum depth of ~800 m and did not intersect a basal contact of the Cerro San Lorenzo Formation. Simplified geological sections, interpreted from data generated through field mapping as well as drill core (Fig. 2.5), suggest a possible thickness of up to 2,500 m. Basalt is variably feldspar‐ (0‐20%) and augite‐ (0‐5%) porphyritic or microporphyritic, and is typically vesicular with 2‐10%, 1‐5 mm amygdules (Fig. 2.7A). Near the B5 deposit (Fig. 2.1), basalt is locally scoria‐like and includes breccia with mm‐scale clasts dominated by bubble wall shards (Fig. 2.7B, C). Massive to pillowed flows are the most common lithofacies and form monotonous nondescript intervals up to several hundred metres thick. Pillows are typically 0.5‐
1.0 m in diameter (Fig. 2.7D), commonly with radial fractures and ‘onion‐skin’‐type concentric flow foliations (Fig. 2.7E). Individual pillow lava flow units are decametres thick and associated Chapter 2 Page 35 with a variety of breccia types (Fig. 2.7F). Pillow fragment breccia is derived from the collapse of pillow mounds that vary from proximal talus breccia to distal facies showing evidence of mass transport, e.g., reverse sorting (Fig. 2.7G). Basaltic autoclastic breccia within the Cerro San Lorenzo Formation includes hyaloclastite and autobreccia. Hyaloclastite has angular and cuspate, mostly pebble‐size breccia clasts with distinct clast outlines and often jigsaw‐fit textures due to quench fragmentation in seawater (Fig. 2.7H). Autobreccia, formed due to viscosity variations within a cooling and flowing lava, display amoeboid‐shaped clasts with abundant, fine (< 1 mm) amygdules (Fig. 2.7I). Autobreccia clasts are typically maroon colour due to hematization and have diffuse clast margins that grade into a darker, chloritic matrix. Volcanic rocks of the Cerro San Lorenzo Formation host all known VMS deposits and prospects (TG1, TG3, and B5). Though outcrop is poor to non‐existent in the vicinity of the deposits, an extensive drill core library was available for study. Construction of a rhyolite‐dacite volcanic complex intimately associated with the VMS deposits is documented by Winter et al. (2004). The extent of the felsic complex is not completely delimited, but at TG1 and TG3 it is a minimum of 2 km in diameter and with a composite thickness of up to at least 300 m. Felsic volcanic rocks represented by massive (Fig. 2.8A) to flow banded lavas, domes or dykes of dacitic to rhyolitic composition and are associated with a variety of volcaniclastic rocks. Breccia is common and varies from units of in‐situ autoclastic breccia (Fig. 2.8B) to transported deposits (Fig. 2.8C). Lava flow‐dome complexes are dominated by buff to light grey, aphyric to weakly feldspar porphyritic, rarely amygdaloidal, but locally spherulitic and perlitic dacite‐
rhyolite. Textures indicate the felsic volcanic rocks are mostly lavas and dykes with associated in‐situ breccias and proximal re‐sedimented volcaniclastic units. Quartz porphyritic lavas are not recognized near the VMS deposits and are generally uncommon throughout the Cerro San Chapter 2 Page 36 Lorenzo Formation. However, quartz‐plagioclase porphyritic rhyolite forms late dykes or stocks within the volcanic complex at TG1 and elsewhere in the formation. Dacitic lava flows and breccias are conspicuous in the immediate hanging wall of the TG1 and TG3 deposits. The dacite is characterized by distinctive pale grey‐green, aphyric to feldspar porphyritic textures, and is commonly amygdule‐rich (Fig. 2.8D). Sedimentary rocks in the Cerro San Lorenzo Formation are limited to thin‐bedded to laminated, black, carbonaceous mudstones in less than 1 m‐thick units. Although volumetrically minor, these pelagic sedimentary rocks are ubiquitous throughout this formation. No other sedimentary rock types are known to occur in the Cerro San Lorenzo Formation. 2.5.2 Cerro El Ereo Formation Though the lower contact with the Cerro San Lorenzo Formation is not exposed, the upper contact with the La Bocana Formation is conformable and relatively sharp. This sequence has a rather limited geographic distribution in the western‐central part of the area and does not extend north of the Las Lomas pluton (Fig. 2.4). This study restricts the rocks included therein to those of similar character that outcrop on the western side of the study area (Figs. 4, 5). The Cerro El Ereo Formation has an approximate thickness of up to 2,000m. An entirely mafic volcanic sequence, the Cerro El Ereo Formation, is defined by distinctive and monotonous coarse ‘crowded’ feldspar porphyritic volcanic and breccia and minor coherent lava. Subhedral feldspar phenocrysts to glomerocrysts range from 1 mm to >10 mm, averaging 4‐5 mm (Fig. 2.9A). Amygdaloidal lava flows are generally not common, though amygdaloidal clasts are locally present in breccia. Volcaniclastic rocks in this formation are typically non‐stratified, matrix‐supported subangular to subround boulder breccia (Fig. 2.9B) or cobble‐ to pebble‐sized lithic and feldspar crystal tuff (Fig. 2.9C). Minor, thin bedded feldspar Chapter 2 Page 37 (and rare pyroxene) crystal and ash tuff, possibly as re‐worked material, occur near the upper contact of the formation (Fig. 2.9D). The sequence is also distinctive in the complete absence of felsic volcanic and sedimentary rocks. No pillow lava or autoclastic breccia similar to those common in the Cerro San Lorenzo Formation are known in the El Ereo Formation. The dominance of poorly vesicular, matrix‐supported, non‐stratified breccia, with distinctive porphyritic juvenile and variably abraded clasts, and the limited geographic extent of the unit, suggests the formation may be the result of a diatreme‐like process. 2.5.3 La Bocana Formation The La Bocana Formation marks a return to bimodal volcanism of basaltic‐andesite and rhyolitic rocks, albeit with a greater abundance of volcaniclastic rocks. The presence of pyroclastic deposits, including crystal‐rich tuff, may indicate a shift to a relatively more shallow water depositional setting. Reyes and Caldas (1987) report an upper Albian age for this sequence based on fossil evidence. The La Bocana Formation has an estimated thickness of 3500 m with conformable upper and lower contacts. Mafic rocks in this unit include highly vesicular, thick flows and dykes (Fig. 2.10A) with well developed flow foliations (Fig. 2.10B). Flows are observed to grade into autoclastic deposits of unsorted coarse breccia (Fig. 2.10C). ‘Breadcrust’ texture is illustrated by chilled and fractured margins of lava flows with interstitial hyaloclastic breccia (Fig. 2.10D). Pillow lava is present in only a few localities. However, polygonal jointed mafic lava flow lobes, apophyses and dykes into volcaniclastic deposits or felsic crystal tuffs (Fig. 2.10E) are common. Felsic rocks range from coherent facies of quartz‐ and/or feldspar‐ porphyritic lava (lava flows, domes, dykes) to crystal‐, lithic‐, and pumice‐bearing tuffs (Figs. 2.10F, G). The La Bocana Formation is the only sequence in the Lancones Basin to include felsic tuffaceous rocks indicative of pyroclastic eruptions. Another enigmatic feature of the felsic volcanic rocks is that Chapter 2 Page 38 granitic xenoliths are present in several flows and were derived either from older crustal basement or from related igneous plutonic roots carried in penecontemporaneous eruptions. Volcaniclastic rocks make up a significant proportion of this unit, and range from chaotic, matrix‐supported, boulder‐size breccia (Fig. 2.10H) to well sorted, pebble‐size breccia and decimeters‐thick cross‐bedded volcanic sandstone (Fig. 2.10I). The boulder breccia occurs as either talus deposits related to fault scarps or mass flows. The well‐sorted, coarse clastic rocks suggest a high energy shallow marine to fluvial environment. Basaltic‐andesite generally dominates the clast composition, though felsic volcanic clasts are common. Locally, calcareous sedimentary clasts are abundant and occur as blocks up to boulder size. Mafic and felsic volcanic rocks are intercalated with the volcaniclastic units and suggest volcanism was also active in a shallow marine setting. 2.5.4 Lancones Formation The Lancones Formation consisting of a basal sequences of polylithic, basaltic‐andesite‐rich, volcaniclastic units represent deposition in a relatively shallow marine environment. These thick‐bedded and variably stratified breccias are intercalated with siliciclastic and calcareous sedimentary units, including limestones, calcareous sandstones, siltstone and greywacke (Fig. 2.11A, B). Sedimentary rocks become more abundant towards the top of the unit. Reyes and Caldas (1987) report fossils in the age range of late Albian and Early Cenomanian. These rocks are considered to be gradational into the forearc sedimentary sequences of the Copa Sombrero Group senso lato in the western Lancones Basin (Fig. 2.4, 11). 2.6 Structural Geology The Lancones basin is situated at a major oroclinal bend along the Andean chain known as Huancabamba deflection (Mégard, 1987), where the Andes trend changes from the north‐
northwest in northern Perú to north‐northeast in Ecuador (Fig. 2.2). Several clockwise rotations Chapter 2 Page 39 are documented north of the Huancabamba deflection, including ~25‐30° and ~25° in Cretaceous‐Tertiary and Tertiary times, respectively (Mourier et al., 1988; Mitouard et al., 1990). The kinematics of the deformation are poorly understood but suggest the opening of the Lancones basin may have been an integral part of the Huancabamba deflection. Part of the rotation of the Amotape block was suggested by Mourier et al. (1988) to have happened in‐situ and related to accretion. Under the model presented (Fig. 2.3C), the opening of the Lancones basin at the northernmost segment of the Andean rift system had a significant component of dextral shear and associated clockwise rotation. Subsequent accretionary tectonism north of the Huancabamba deflection account for the additional rotations. (Fig. 2.3 D‐E). Under this scenario, the Albian rift segment at the Lancones basin may not have differed substantially from that of the WPT with respect to an original NNW‐trending geometry. Several deformation events are recognized in northwestern Perú in Late Cretaceous and Tertiary times (Jaillard et al., 1999), including a mid‐ to late‐Albian tectonic phase correlated with the Mochica orogeny defined in the central‐northern Andes (Mégard, 1984). This deformation is manifested mostly as broad open folds in the eastern Lancones basin, in general with northeast‐southwest striking, ‘Andean‐normal’, fold axes (Figs. 4, 5). A series of WNW‐
trending topographic lineaments are identifiable from the satellite data, but do not appear to have resulted in any mapable displacement (Fig. 2.4). These linears may be representative of the same fracture system that controlled the emplacement of the plutonic rocks. Data derived from oil exploration within the basin (Fig. 2.1) provides additional constraints on the Mesozoic structural history of the area. Alencastre (1980) reported a number of faults, interpreted from petroleum borehole and geophysical data to be as normal or block and transcurrent faults. Seismically imaged faults in the Tertiary as well as Cretaceous strata strike Chapter 2 Page 40 east‐northeast to northeast. Tegart et al. (2000) suggested northeast‐trending subsurface Paleozoic faults imaged to the southwest under cover of the Sechura Basin (Alencastre, 1980) can be projected to the Lancones Basin and inferred these to be primary graben‐bounding faults controlling the location of Cretaceous volcanic rocks and VMS deposits. 2.7 Plutonic Rocks As with the segment of the Coastal batholith within Western Peruvian Trough, the batholith in northwestern Perú has been emplaced within the marginal volcanic arc succession of the Lancones basin (Fig. 2.1). Limited plutonism has occurred in the Copa Sombrero Group to the west. Although much of the Tambogrande district has been intruded by various phases of the Coastal batholith, more voluminous plutonism is known to the east of the map area (Fig. 2.4). The Las Lomas complex, a 15 km‐wide zoned gabbroic to granitic intrusion in the center of the map area, yields U‐Pb and Ar‐Ar ages that suggest the time of emplacement for these rocks was 47‐88 Ma (Chapter 6). No outcropping intrusive phases have been identified as syn‐volcanic in origin. As the granitoid suites occur within the northeast trending Lancones basin, the Coastal batholith of northwestern Perú differs in trend from the northwest‐trending central Peruvian batholith (Fig. 2.4). Locally, the linear contacts of the granitoid suites (e.g., Las Lomas Complex) also mimic the northeast trend. The northeast ‘Andean’ trend in this region is thus characteristic of both the Coastal batholith and the host volcanic arc suites. However, as the Las Lomas complex has been rotated ~25° clockwise (Mitouard et al., 1990), the original orientation may have been closer to north‐trending. Chapter 2 Page 41 2.8 U‐Pb Geochronologic Data Sample preparation and analytical procedures for U‐Pb geochronology are described in Appendix A. Two U‐Pb techniques were utilized because of the subtle inherited zircon cores to many of the zircon populations. Thermal Ionization Mass Spectrometry (TIMS) and Sensitive High Resolution Ion Microprobe ‐ Reverse Geometry (SHRIMP‐RG) analytical data are presented in Tables A1‐A2. Rock sample and zircon descriptions are provided in Table A3 and shown on the map (Fig. 2.5). A total of thirteen U‐Pb zircon analyses are discussed in detail below. Summarized results are also plotted on schematic regional‐ and mine‐scale stratigraphic sections (Fig. 2.12). All errors for these ages are presented as 2 sigma unless otherwise specified. Felsic and intermediate volcanic rocks sampled from the project area are generally aphyric or weakly porphyritic and were typically found to yield low quantities of zircon concentrates. In addition, some samples analyzed by TIMS produced an older age due to an inherited Proterozoic component. Therefore, a combination of TIMS and SHRIMP‐RG analysis was selected due to the variability of zircon contents as well as to characterize igneous crystallization and inherited components. Amongst the samples processed, with two exceptions, only quartz porphyritic varieties of felsic volcanic rocks contain zircon. Nine rock samples of dacitic to rhyolitic volcanic rocks from the eastern Lancones basin were dated by U‐
Pb zircon methods, including three samples from the Cerro San Lorenzo Formation and six samples from the La Bocana Formation. An additional four samples from the La Bocana Formation displayed strong inheritance and did not yield igneous ages. U‐Pb zircon ages are presented from four TIMS and six SHRIMP‐RG analyses. An additional ~twenty‐five samples were processed but did not yield zircons. Chapter 2 Page 42 2.8.1 Volcanic Rocks of the Cerro San Lorenzo Formation Sample TG1‐136 (rhyolitic volcaniclastic) represents the immediate hanging wall strata to the TG1 massive sulphide deposit. This and all other samples yielded zircons that were mostly clear, colorless, stubby to elongate, euhedral prisms. The sample produced a small yield of <74 µm zircons that appear entirely magmatic. Twelve grains were analyzed by SHRIMP‐RG of which three data points cluster at slightly lower ages and one slightly higher than the main population and these were omitted from the age calculation. The weighted mean 206Pb/238U age of eight analyses is 104.8 ± 1.3 Ma with MSWD of 1.8 (Fig. 2.13A). Sample LW‐016 (quartz porphyritic rhyolite dyke) was also dated by SHRIMP‐RG analysis. All crystals appear to be devoid of xenocrysts and display well developed oscillatory to lamellar zoning under cathodoluminescence (CL; Appendix A). Excluding two samples that have slightly higher 206Pb/238U ages than the main population, ten analyses of both cores and rims yield a single population with a weighted mean 206Pb/238U age of 103.2 ± 1.0 with MSWD of 1.3 (Fig. 2.13B). Sample TG1‐111 (quartz‐feldspar porphyritic rhyolite dyke cutting the TG1 massive sulphide deposit) yielded five zircon fractions. Fractions A and E yield overlapping concordant analyses (Fig. 2.14A) with a total range of 206Pb/238U ages of 100.2 ± 0.5 Ma, which is interpreted to be the best estimate for the crystallization age of the sample. Fraction B is also concordant but yields a slightly younger 206Pb/238U age, and has possibly suffered minor post‐crystallization Pb‐
loss. Fractions C and D yield discordant analyses with older 207Pb/206Pb ages, indicating the presence of minor older inherited zircon components in some of the grains, presumably as “cryptic” cores that could not be distinguished visually. Chapter 2 Page 43 2.8.2 Volcanic Rocks of the Lancones Formation A moderate amount of zircons produced from sample LW‐086 (a single clast from a rhyolite boulder breccia) was dated using TIMS. Six fractions were analyzed (Fig. 2.14B). Fractions B and C yield overlapping concordant analyses with a total range of 206Pb/238U ages of 99.3 ± 0.3 Ma, which is interpreted to be the best estimate for the crystallization age of the sample. Fractions E and F have suffered minor post‐crystallization Pb‐loss, and fractions A and D gave discordant analyses with older 207Pb/206Pb ages, indicating the presence of minor older inherited zircon components in these fractions. Sample LW‐013 is a quartz porphyritic rhyolite dyke within the upper Cerro San Lorenzo Formation near the contact with the La Bocana Formation. A small quantity of zircons was recovered. CL images indicate that a few crystals have small xenocrystic cores, though most display simple oscillatory zoning (Appendix A). SHRIMP‐RG analyses of twelve crystals yield no systematic measurable difference in ages between cores and rims of the wholly magmatic crystals but show some scatter between 206Pb/238U ages of 96.6 and 104.6 Ma. Three analyses with slightly higher 206Pb/238U ages than the main population were rejected and are likely xenocrystic. As there is no geologic reason to reject any of the remaining zircon U‐Pb ages, a weighted average 206Pb/238U age of 99.8 ± 1.6 with MSWD of 3.2 is calculated (Fig. 2.13C). The high MSWD reflects the scatter of the data. Most of the zircons from sample LW‐078 (quartz‐feldspar porphyritic clast in polylithic breccia) have CL patterns that suggest entirely magmatic phases with oscillatory to lamellar zoning in igneous cores, though several grains have inherited zircon cores (Appendix A). Twelve crystals were analyzed by SHRIMP‐RG. Sample point 1 yielded a slightly younger age (95.4 Ma) than the main trend, possible due to Pb‐loss though it was an analysis of a zircon crystal core, Chapter 2 Page 44 and was omitted from the age calculation. Eleven data points provide a 206Pb/238U age of 98.9 ± 0.7 Ma with MSWD of 1.1 (Fig. 2.13D). Four zircon fractions from sample LW‐010 (dacite feldspar porphyry) were analyzed by TIMS and all yielded concordant analyses (Fig. 2.14C). Fractions A, C and D yield overlapping error ellipses with a total range of 206Pb/238U ages of 97.0 ± 0.5 Ma, which is interpreted to be the best estimate for the crystallization age of the sample. Fraction E yields a slightly younger 206
Pb/238U age, indicating that this fraction suffered minor post‐crystallization Pb‐loss. Sample LW‐077 (dacite flow and autobreccia) contained abundant zircons and was analyzed by TIMS with six fractions all giving concordant analyses (Fig. 2.14D). The analyses scatter along Concordia, however, likely due to variable effects of post‐crystallization Pb‐loss that were not completely eliminated by strong air abrasion. Assuming that no older inherited zircon cores are present in any of the grains, a minimum crystallization age for the sample is given by the oldest 206
Pb/238U age (for fraction B) at 95.0 ± 0.5 Ma. Most zircons from sample LW‐043 (rhyolite quartz‐feldspar porphyritic flow) in CL imagery display inherited cores represented by dark, resorbed zones with oscillatory zoned overgrowths (Appendix A). The sample was dated using SHRIMP‐RG with nearly all data points selected from crystal rims. Two of 12 data points yield substantially younger ages (due to possible significant Pb‐loss?) and one additional data point with slightly older 206Pb/238U age (95.2 Ma), due to distinctly different and high U contents (1586 ppm), were rejected. The weighted mean 206
Pb/238U age for nine data points is 91.1 ± 1.0 with a MSWD of 3.7 (Fig. 2.13E). A number of additional samples display strong Proterozoic inheritance and a meaningful igneous age could not be calculated. Sample LW‐051 (quartz‐lithic rhyolite tuff) yielded a sufficient amount of zircon for three fractions which were analyzed by TIMS (Fig. 2.14E). Chapter 2 Page 45 Fractions A and B consisted of elongate prismatic grains, whereas C and D were stubby prisms. Fraction B is slightly discordant near ~70 Ma on the concordia curve; however fractions A and C are moderately to strongly discordant, indicating that at least these two fractions contained a significant component of older inherited zircon cores. Two‐point discordia lines through fraction B and each of the two more discordant analyses yield calculated upper intercept ages of 1.16 Ga and 1.51 Ga, suggesting a mainly Mesoproterozoic average for the inherited zircon component. The 206Pb/238U age of fraction B is surprisingly young (73.5 Ma), however, indicating either that this sample has a considerably younger crystallization age than all of the other dated samples, or that this fraction experienced much stronger post‐crystallization Pb‐
loss than was evident from the systematics of any of the other samples. A crystallization age for this sample cannot be assigned on the basis of the available analytical data. Sample LW‐066 (quartz porphyritic rhyolite breccia) yielded a small quantity of zircon analyzed by SHRIMP‐RG. Most samples show strong Pb‐loss and unrealistically young 206
Pb/238U ages whereas older ages were measured ranging to 1898 Ma. Two data points yield 206
Pb/238U ages of 99.4 and 99.8 Ma are in agreement with the range of ages determined for other volcanic rocks in this area (97.0 to 99.3 Ma). Zircons from samples LW‐026 (quartz‐feldspar porphyritic rhyolite dyke) and LW‐033 (quartz‐feldspar porphyritic rhyolite breccia) also yield dominantly older and inherited 206
Pb/238U ages ranging up to ~2540 Ma. Single zircons from each of LW‐026 and LW‐033 yield 206
Pb/238U ages of 90.1 and 90.9 Ma, respectively. These ages are in agreement with ca. 91 Ma rhyolite (LW‐043) sampled in this part of the Lancones Formation. Chapter 2 Page 46 2.9 Discussion 2.9.1 Depositional Evolution of the Lancones Basin The ~10 km thick stratigraphic package that defines the volcanic arc sequence of the Lancones Basin represents a large bimodal volcanic eruptive event in mid‐Cretaceous times within a marginal basin that presently is 150 km wide. The stratigraphy suggests a progressively shallowing basin and evolving depositional environment through mid‐Cretaceous times. Although the continental crust at the Pacific margin would have been significantly thinner than the present day western Andes, the amount of crustal attenuation and subsidence was substantial. An ensialic rift event related to a ‘Mariana‐type’ supra‐subduction zone is envisaged to account for the volcanism. The total inferred thickness of the Lancones basin successions implies a syn‐subsidence depositional environment. Rifting, crustal subsidence, volcanism and sedimentation were synchronous. The Cerro San Lorenzo Formation is interpreted as a relatively deep water facies based on the absence of pyroclastic rocks and presence of pelagic sedimentary rocks. Although not unequivocal evidence of paleowater depth, a volcanic succession comprised entirely of massive to pillowed lava flows and associated autoclastic deposits may be interpreted as a result of deep water suppression of erupting lavas resulting in effusive‐only eruptions (Cas, 1992; Batiza and White, 2000; White et al., 2003). A deep water environment is also consistent with the inferred depositional setting and deposit style (i.e., metal budget, alteration) of VMS deposits at Tambogrande (Tegart et al., 2000; Winter et al., 2004). The Cerro San Lorenzo Formation lacks abundant siliclastic or distal volcaniclastic deposits despite the inferred deep water setting in a narrow rifted continental basin. With continental topographic highs inferred at both margins, a greater proportion of sedimentary rocks might be expected in a deep basin, or at least close to the margins of the basin. Indeed, the San Pedro Chapter 2 Page 47 Group, interpreted as Albian in its upper parts, has been mapped along the eastern portions of the Lancones Basin (Reyes and Caldas, 1987) and may be the slope facies of the eastern margin of the basin in the early‐syn rift phase. The Copa Sombrero Group forearc sedimentary sequence temporally overlaps with the entire volcanic arc succession along the western margin. Clearly sedimentation was active during the volcanic episodes. Therefore, it is possible that the Cerro San Lorenzo Formation represents an ocean ridge‐type structure, analogous to back‐arc volcanic ridge edifices in modern settings. For comparison, the Manus Basin continental back‐
arc rift hosts an active 700 m‐high bimodal volcanic ridge flanked by siliciclastic rocks (Taylor et al., 1995). If the Cerro San Lorenzo Formation had been built as a topographically high volcanic edifice, the interstratification of volcanic/volcaniclastic and sedimentary rocks may be limited to the flanks of the volcanic ridge. All known VMS deposits and prospects are associated with the Cerro San Lorenzo Formation, the lowermost of the volcanic sequences of the Lancones Basin. The Cerro San Lorenzo Formation represents the initial magmatic pulse, or phase 1 volcanic event, of the rift sequence and appears to have been deposited in a relatively deep water environment during a period of maximum subsidence. During phase 2 volcanism, successions become progressively more volcaniclastic‐ and sediment‐rich and appear to have been deposited in relatively shallow water as late‐rift, basin‐fill sequences. Therefore, the association of VMS deposits with the Cerro San Lorenzo Formation may be linked to the establishment of large‐scale hydrothermal systems in the early‐rift tectonic stages, during the most likely period of high magma input, maximum heat flow and high permeability due to rift‐related extensional and transcurrent faulting. The phase 2 volcanic sequences have less potential for the discovery of VMS deposits and are not known to host any VMS‐type mineral occurrences. Chapter 2 Page 48 2.9.2 Timing and Duration of the Volcanic Arc U‐Pb TIMS and SHRIMP‐RG zircon ages presented in this paper permit an estimation of the timing and duration of the arc volcanism in the Lancones Basin (Fig. 2.12). Volcanism commenced at a minimum of ca. 105 Ma (Albian) and continued to at least ca. 91 Ma (Turonian) suggesting a minimum volcanic arc lifespan of ~14 Ma. The upper age limit of the volcanic succession is further constrained with U‐Pb zircon ages of 78‐88 Ma for granitic rocks of the Coastal batholith that intruded into Lancones basin. In the Ecuadorian segment of the Lancones basin, granodiorite at the Los Linderos porphyry Cu‐Mo‐Au prospect (Chiaradia et al., 2004) cuts basaltic rocks and is dated at 88.4 ± 1.0 Ma (Appendix B). Compared to the youngest Lancones basin volcanic rocks dated in this study (91.1 ± 1.0 Ma), there is a minimal temporal transition from arc volcanism to arc plutonism. Likewise, a temporal overlap from volcanism to plutonism is recognized within Coastal Batholith the Western Peruvian Trough (Soler and Bonhomme, 1990). The duration of volcanic phases is also well constrained based on the data reported herein (Fig. 2.12). Phase 1 ranges from an oldest age of 104.8 ± 1.3 Ma (Sample LW‐16) to a minimum of 100.2 ± 0.5 Ma (sample TG1‐111) indicating a minimum lifespan of ~4‐6 Ma. Phase 2, limited to dates from the La Bocana Formation, ranges from a maximum age of 99.8 ± 1.6 Ma (sample LW‐086) to a lowermost age of 91.1 ± 1.0 Ma (LW‐043) and suggests a minimum duration for volcanism of 7‐10 Ma. No volcanic rocks from the El Ereo Formation were suitable for dating methods employed in this study, however, the approximate age range can be deduced from the given ages of the overlying La Bocana Formation and underlying Cerro San Lorenzo Formation (Fig. 2.15). Volcanic rocks of the El Ereo Formation must be younger than 100.2 Ma. However, the minimum age requires some consideration of the stratigraphy and spatial distribution of the Chapter 2 Page 49 various formations. Firstly, the oldest rocks in the La Bocana Formation occur in the northern part of the map area and are dated at 99.3 ± 0.3 Ma. These values suggest a potential 0‐3 Ma interval for eruptions of the El Ereo Formation lavas. However, the contact between the Cerro San Lorenzo Formation and La Bocana Formation is uncertain and the El Ereo Formation is not present in the north part of the map area (Figs. 2.5, 2.14). Conversely, south of the Las Lomas complex, the La Bocana Formation is known to be conformable above the El Ereo Formation. Age constraints in this part of the La Bocana Formation are much younger (91.1 ± 1.0 Ma) indicating a time interval than spans at least ~8‐11 Ma for duration of the Cerro San Lorenzo and La Bocana formations. Under this scenario in the southern Lancones Basin, the El Ereo Formation would have potentially been deposited contemporaneously with the deposition of the lower part of the La Bocana Formation to the north. Due to the monotonous basaltic volcanism and paucity of sedimentary rocks in the El Ereo Formation, a ‘single’ large volcanic event is postulated that may not have required a significant time interval. This is consistent with a possible diatreme scenario and a maar‐like eruption. 2.9.3 Age of Massive Sulphide Deposits Winter et al. (2004) provide a detailed reconstruction for the TG1 and TG3 VMS deposits whereby equivalent time‐stratigraphic horizons can be correlated between each of the TG1 and TG3 deposits. Therefore, the relative timing of massive sulphide formation for these deposits is considered practically identical. The B5 VMS deposit ~10 km to the south could not be represented on a continuous section with TG1 and TG3 but is hosted in similar volcanic strata. Moreover, lead isotopic signatures (Chapter 5) are identical for all three deposits and suggest a genetic, and possible temporal, relationship. The minimum age of massive sulphide formation at TG1 is well constrained by several U‐Pb zircon dates. The immediate hanging wall felsic volcaniclastic rocks have been dated at 104.4 ± Chapter 2 Page 50 1.9 Ma (sample TG1‐136). The sample is a sand‐ to pebble‐size, moderately‐well sorted, dacitic volcaniclastic unit that, based on the volcanic reconstructions of the depositional environment (Chap. 3), represents a distal facies of a seafloor rhyolite flow dome complex. The rhyolite eruption was most likely synchronous with or slightly post‐dates massive sulphide formation. Additional age constraints on the deposits are thus provided by a rhyolite dyke that intrudes the TG1 deposit. This felsic unit represents a post‐mineralization phase and provides a minimum age limit of 100.2 ± 0.5 Ma for formation of the TG1 and TG3 deposits. Footwall rocks to the VMS deposits did not yield zircons or other minerals suitable for U‐Pb zircon dating, and therefore the maximum age for the massive sulphide system could not be determined. 2.9.4 Comparison of the Lancones Basin to the Western Peruvian Trough Geochronologic and stratigraphic data presented herein verify the broadly contemporaneous evolution of the volcanic successions of the Lancones basin and Western Peruvian Trough. Stratigraphically, the Casma group that filled the Western Peruvian Trough has a total thickness of up to 9 km and is dominated by mafic with lesser felsic volcanic rocks (Myers, 1974; Cobbing et al., 1981; Offler et al., 1980; Soler and Bonhomme, 1990), comparable to the 8‐10 km of mafic‐dominated bimodal strata of the Lancones Basin (Fig. 2.16). Volcanism recorded by the Casma Group was largely terminated during the late Albian‐Early Cenomanian indicating a short‐lived but vigorous volcanic event. The Mochica phase, which represents the first contractional orogeny in the central Peruvian Andes (Mégard, 1984), affected the Casma Group at the Albian‐Cenomanian boundary. This effectively represents the termination of the rift‐stage marine volcanism. Volcanic rocks post‐
dating the Mochica phase were interpreted to be as the products of shallow marine to subaerial volcanism (Webb in Cobbing et al., 1981) and these eruptions overlapped the initial phases of the Coastal Batholith (Soler and Bonhomme, 1990). A marine transgression and the deposition Chapter 2 Page 51 of shelf carbonates in the Western Peruvian Trough are recorded during Cenomanian to Turonian times (Jaillard and Soler, 1996). Although deformation related to the Mochica orogeny is not well documented in the volcanic rocks of the Lancones Basin, the timing of the contractional episode coincides with the transition from the Cerro San Lorenzo Formation to La Bocana Formation. Moreover, the facies change from deep(?) water, pillow lava‐dominated volcanic rocks of the Cerro San Lorenzo Formation to mixed pyroclastic, siliciclastic and carbonate rocks, representing dominantly shallow water environments, of the La Bocana Formation and younger sequences suggests a major geodynamic change, the timing of which is similar to that described for the upper Casma Group. The stratigraphic thickness, relative age, volcanic facies and inferred depositional environment of the Western Peruvian Trough is remarkably similar to the Lancones Basin and suggests that both basins evolved along nearly identical pathways. 2.9.5 Tectonic Implications The fundamental processes required to generate a marginal rift basin related to subduction along a continental margin are considered to be fairly well understood with the prerequisites being a subducting and ‘sinking’ ocean slab and a trench line that migrates oceanwards at a greater rate than the overriding plate, termed slab ‘rollback’ (Hamilton, 1995). Benavides‐
Cáceres (1999) suggests the Lancones basin marginal rift resembled a ‘Mariana‐type’ arc (Karig et al., 1978) which forms when the overriding plate, either oceanic or continental, is attenuated and rifted due to tensional stresses associated with a steeply dipping slab, rollback and arc migration. The stratigraphic and geochronologic data in this study help to reconcile the tectonic regime related to the formation of the Lancones basin and, by inference, the Western Peruvian Trough. Chapter 2 Page 52 The formation of the Western Peruvian Trough is temporally linked with a major global tectonic event, the final break‐up of Gondwana, which culminated in the opening of the South Atlantic Ocean in the Early Cretaceous (Scotese, 1991). More specifically, the Albian represents the time of final separation of South America from Africa and opening of the equatorial South Atlantic ca. 105 Ma (Sibuet et al., 1984; Scotese, 1991). The lack of significant westward motion of the South America plate during the Albian was of major importance to the formation of the Western Peruvian Trough and Lancones basin as a steeply subducting but retreating oceanic slab along the Pacific margin of South America triggered oceanwards (westwards) migration of the overlying volcanic arc (Soler and Bonhomme, 1990). The result was extension and rifting of the overlying western continental margin of South America and subsequent arc volcanism leading to a paleogeography that may have resembled the modern day Bransfield Strait (Barker and Austin, 1998) or the Miocene Japan Sea (Jolivet and Tamaki, 1992). With the onset of spreading in the South Atlantic and the westward drift of the South American continent, the development of the Lancones Basin and Western Peruvian Trough terminated (Soler and Bonhomme, 1990). This transition from extensional to contractional tectonics is marked by the Mochica Orogeny (Mégard, 1984). In the Lancones Basin, shallow marine to terrestrial volcanism of phase 2 followed in a less extended margin and evolved to granitoid plutonism under a continental arc regime. 2.9.6 Inheritance in Zircons and Implications for Basement Rocks Only minor inheritance is suggested from the zircon data from rocks of the Cerro San Lorenzo Formation but significantly more inheritance is obvious from volcanic rocks from the La Bocana Formation. Although there is somewhat of a sampling bias with more samples collected from the La Bocana Formation than from the Cerro San Lorenzo Formation, the greater inheritance of zircon by felsic rocks of the La Bocana Formation may indicate a more important Chapter 2 Page 53 continental crust component in the source and/or pathways of these lavas. Three of ten samples from the La Bocana Formation display a variety of inherited ages up to 2,500 Ma. The strong inheritance implicates continental crust in the genesis of these rocks, especially in the younger sequences. This concept has implications for basement development and magma genesis and is explored in greater detail elsewhere (Chapter 5). 2.10 Conclusions The depositional history of the Lancones Basin is recorded by a ~10 km sequence of bimodal volcanic and volcaniclastic rocks representing at least ~14 Ma of volcanic activity from ca. 105 Ma to ca. 91 Ma based on U‐Pb zircon ages. Two main phases of volcanism comprise the volcanic arc sequence: (i) Phase 1 volcanism, recorded by the Cerro San Lorenzo Formation, represents deep‐water, bimodal, mafic‐dominated volcanic rocks comprised of lava flows, autoclastic breccias and minor pelagic sedimentary rocks. Four U‐Pb zircon ages ranging from 104.8 ±1.3 to 100.2 ±0.5 Ma suggest middle to late Albian ages for this formation. VMS deposits at Tambogrande are hosted within the Cerro San Lorenzo Formation and have a minimum age of 104.8 ±1.3 Ma, as constrained by the pre‐syn mineralization felsic volcanic rocks. Pre‐mineralization volcanic rocks did not yield any minerals suitable to U‐Pb dating methods. (ii) Phase 2 volcanism is recorded by the Cerro El Ereo, La Bocana and Lancones Formations. These formations are dominated by deposits of shallow marine facies and are represented by bimodal volcanism. Five U‐Pb zircon ages range from 99.3 ±0.3 to 91.1 ±1.0 Ma. The volcanic phases are broadly chronologically correlated with major tectonomagmatic events, specifically the opening of the South Atlantic Ocean in the Early Cretaceous during the final break‐up of Gondwana. Phase 1 volcanism was a result of strong crustal attenuation and rifting of the western margin of a relatively static South American continent during Albian Chapter 2 Page 54 times. Phase 1 correlates with pre‐opening or rift stage of the South Atlantic at equatorial latitudes. The Lancones basin during phase 1 volcanism would have resembled a modern day back‐arc basin or rifted arc‐type setting. Active spreading in the South Atlantic Ocean and westward drift of the South American continent correlates broadly with the Mochica orogeny in the western Andes during Albian times. This contractional tectonic event marks the termination of the Lancones basin phase 1 volcanic event. Phase 2 volcanism is more typical of Andean arc magmatism and was the result of subduction beneath a less extensional margin during late Albian to Turonian times. Chapter 2 Page 55 Figure 2.1. Morphostructural units of the Peruvian Andes (modified after Benavides‐Cáceres, 1999). Cretaceous marginal basins ‐ Lancones (LB), Huarmey (HB) and Cañete (CB) basins ‐ are superimposed. Also shown are the locations of VMS deposits and prospects (circles) (data from Steinmüller et al., 2000). Chapter 2 Page 56 Figure 2.2. A. Location map for the Tambogrande project; B. regional map showing major tectonostratigraphic units of coastal northwestern Perú. The locations of VMS deposits (TG1, TG3, and B5) in the Tambogrande area are also shown and field area of this study outlined. Modified after Jaillard et al. (1999) and Tegart et al. (2000). Chapter 2 Page 57 Table 2.1. Summary of location and description data for samples analyzed in this study for U‐Pb zircon dating. Calculated UTM UTM age +/‐ Description Sample ID Easting Westing error (Ma) Analytical method Comments Cerro San Lorenzo Formation TG1‐136 573712 9454484 104.4 +/‐1.9 Rhyolitic volcaniclastic (TG1 deposit hanging SHRIMP‐
wall); diamond drill hole 00TG1‐136 @ 50.20 RG metres; poor‐moderately sorted, sand‐pebble sized, re‐sedimented autoclastic breccia. LW‐016 588523 9480533 104.7 +/‐1.4 Rhyolite porphyry: dyke with very fine grained SHRIMP‐
quartz phenocrysts. RG TG1‐111 573696 9454532 100.2 +/‐0.5 Rhyolite porphyry breccia: diamond drill hole 99‐ TIMS
TG1‐111 @ ~78‐88 m; massive, light grey to buff, feldspar (5‐10%, 3‐4 mm) and (<3%, ~1 mm) phenocrysts. La Bocana Formation LW‐086 595654 9496749 99.3 +/‐ 0.3 Rhyolite breccia: 1.5 m thick debris flow unit TIMS
within a ~10 m thick volcaniclastic sequence above mafic volcanic rocks; sample is 1 ~0.5 m clast of buff colour felsic lava with feldspar (5%, 3‐4 mm) and quartz (2‐3%, <2 mm) phenocrysts.
LW‐013 LW‐078 LW‐010 LW‐077 588255 9482680 99.1 +/‐ 1.4 Rhyolite porphyry: quartz phenocrysts; columnar SHRIMP‐
joints. RG 589870 9490590 98.8 +/‐1.0 Polylithic boulder breccia: dominated by SHRIMP‐
basaltic‐andesite, lesser rhyolite and minor RG calcareous sedimentary clasts; sample is >1 m, subround, feldspar‐quartz porphyritic clast; 5‐7% pink, stubby feldspars to 5 mm; 5‐7% quartz to 5‐6 mm. 587799 9489045 97.0 +/‐0.4 Dacite porphyry: light green, aphanitic, aphyric, TIMS
stubby plagioclase (5%, 1‐2 mm) phenocrysts, massive, partly sericitized. 580452 9501568 90.3 ‐ 95.3 Dacite: flow and breccia with partially digested TIMS
granitic xenoliths. Chapter 3 only ~ 30 zircons recovered, nearly all < 74 µ
12 grains analyzed.
A small amount of clear, colorless, stubby prismatic zircon recovered. All zircons >74 µ in diameter were picked and subjected to strong air abrasion, and then split into five sub‐equal fractions A, B, C, D, and E. 4 fractions: A, B, C) 1 fraction with 3 splits; slightly yellowish stubby prisms with l/w of 2‐3:1; less 10% broken crystals; D) 14 crystals >104 µ; colourless and clear, broken, anhedral, +/‐ solid inclusions, mostly fragments of prismatic crystals; E) 17 crystals >>74 µ; slightly yellow to clear with minor inclusions, l/w = 3‐4:1; euhedral, 10% broken; F) 32 >74 µ; clear needles with rare solid inclusions; euhedral, l/w=4‐6:1; 10% broken. recovered only small yield; 12 zircons analyzed.
recovered only small yield; 12 zircons analyzed.
5 fractions/subfractions: A) 42 grains 74‐104 µ; B,C,D,E) ~250 grains <74 µ; 4 splits; stubby prisms l/w=2‐3:1. Picked 95% of zircons recovered. Analyzed nearly all zircons recovered (~300). 95% < 74 µ, clear to pink, stubby prisms; 6 splits. Page 58 Table 2.1 continued. Calculated Description UTM UTM age +/‐ Analytical Sample ID Easting Westing error (Ma) method Comments LW‐043 567138 9482242 91.1 +/‐1.0 Rhyolite porphyry: flow‐banded dyke; quartz and SHRIMP‐ recovered only small yield; 12 zircons analyzed.
feldspar phenocrysts. RG LW‐051 562449 9479084 Proterozoic Rhyolite tuff: quartz crystal‐rich, lithic‐bearing; TIMS
inheritance quartz to 5‐6 mm. LW‐066 588107 9489337 ~99.8? Rhyolite porphyry breccia: rhyolite breccia mixed SHRIMP‐
Proterozoic with sedimentary rock clasts; sample is from 20 RG inheritance separate <20 cm subround clasts; quartz (5%, 1‐
2mm) phenocrysts. 569870 9478046 Proterozoic Rhyolite porphyry: dyke; light grey‐green; quartz SHRIMP‐
inheritance (<5%, < 1mm) and feldspar (5%, 1 mm) RG phenocrysts. LW‐026 LW‐033 569685 9478490 Proterozoic Rhyolite porphyry breccia: polylithic, mafic‐
SHRIMP‐
inheritance dominated volcanic breccia; sample from a ~70 RG cm rhyolite clast with feldspar‐quartz phenocrysts. 4 fractions: A) >104 µ; l/w=5/1; colourless, clear with complex terminations; B) >74 µ, 1 crystal >104; stubby and broken crystals; l/w = 2‐3/1; possible zoning; C,D) small, <74 µ, stubby prisms, clear to pink, l/w=~2:1. Picked nearly all zircons recovered. 18 zircons recovered; 11 grains analyzed, all < 104. Mostly stubby prisms, colourless to slightly pink. 17 grains recovered; 11 grains < 104 µ; some broken crystal fragments in the >104. recovered only small yield; 13 data points; 12 grains.
Chapter 3 Page 59 Figure 2.3 (on the following page). Schematic paleogeographic model of the development of Perú‐Ecuador segment of the western margin of South America (SA) from the Jurassic to present using data from Mourier et al. (1988), Mitouard et al. (1990), Litherland et al. (1994), Aspden et al. (1995), Noble et al. (1997), Arculus et al. (1999), Benavides‐Cáceres (1999), Jaillard et al. (2000), Bosch et al. (2002), and Polliand et al. (2005). A. Jurassic to earliest Early Cretaceous: ~SE‐directed convergence of the proto‐Farallon‐Caribbean ocean plate with continental SA; subduction occurs along the Ecuadorian segment, whereas the Peruvian NNW‐
trending margin is a sinistral transform; Amotape terrane is a micro‐continent approaching SA; B) change in convergence direction from SE to ~NE; accretion of the Amotape terrane, notably along the Peruvian segment; dextral faulting of Amotape terrane and clockwise rotation of blocks; ocean‐continental plate boundary ‘jumps’ toward the west; during this period the NW‐
trending Peruvian margin becomes a subduction zone whereas the Ecuadorian NE‐trending margin becomes a dextral transform. C) trench ‘roll‐back’ occurring along Peruvian margin and extension in overriding SA plate; the Lancones basin and Western Peruvian Trough open up along a margin parallel rift and result in the deposition of Cretaceous sedimentary and arc volcanic rocks; continued dextral displacement of Amotape terrane; D) termination of marginal rifting, accretion of ocean plateau ‘Pallatanga’ terrane in Ecuador, and deformation of Andean terranes; formation of Macuchi island arc near margin to be accreted to Ecuadorian segment by Early Oligocene; E) modern day tectonostratigraphic model; compressive tectonic regime; E‐
directed convergence. Chapter 2 Page 60 Chapter 2 Page 61 Figure 2.4. Regional geologic map for the portion of the Lancones basin reviewed in this study; modified from Reyes and Caldas (1987) and from mapping during this study. The location of samples for U‐Pb zircon geochronologic studies are shown and labeled by sample name. Lines A‐A’ and B‐B’ show the trace of sections in figures 2.5 a‐b. Chapter 2 Page 62 Figure 2.5. A. Regional geological cross section A‐A’ through the southern region of the map area. Looking northeast. TG1 and TG3 massive sulphide deposits projected from the south. B. Regional geological cross section B‐B’ through the northern region of the map area. Looking northeast. Legend as per figure 2.4. See map in Figure 2.4 for trace of sections. Chapter 2 Page 63 Figure 2.6. Schematic stratigraphic column of the eastern portion of the Lancones basin. Inset section shows the Tambogrande area in more detail. Chapter 2 Page 64 Figure 2.7 (following page). Field and drill core photographs of mafic rocks from the Cerro San Lorenzo Formation: A. Feldspar porphyritic and amygdaloidal basalt. B. Drillcore from the B5 area, aphyric basalt with autobrecciated margin and close up of breccia C. Illustrates the highly vesicular, scoria‐like clasts; note the small fragments of bubble‐wall shards. D. Section through basaltic pillow lavas at Rio Quiroz; pillows are up to 1 m wide; individual pillow flows are up to 10’s of metres thick. E. Basaltic pillow lavas displaying well developed concentric flow foliations; this specimen is partly broken along radial fractures. F. Pillow basalt unit (P) overlain by mass flow (MF) deposit of basaltic pillow lava and autobreccia clasts. A mafic dyke (Dk) cuts both units and would have probably supplied lava to another cycle of pillow lavas and breccias. G. Medium to thick bedded basaltic volcaniclastic deposits ranging from sand‐ to boulder‐size; note the reverse sorting of the thicker (~1m) basal unit (see arrow) possibly indicative of a massive flow. H. In‐situ autoclastic (hyaloclastic) breccia. Note the jigsaw‐fit textures of the clasts. Breccia are gradational into massive lavas. Drill core, B5. I. In‐situ autoclastic breccias from drillcore, TG3. Bulbous‐shaped clast with diffuse margins in a dark green chlorite matrix; margins of clasts display fine (sub mm) chloritic amygdules. This breccia grades into massive lava. Chapter 2 Page 65 Chapter 2 Page 66 Figure 2.8. Drill core photographs of intermediate and felsic rocks from the Cerro San Lorenzo Formation: A. Massive feldspar porphyritic rhyolite. Scale units are mm. TG1 area. B. Flow‐
banded rhyolite autobreccia; these breccias typically grade into lavas. Textures partly masked by quartz and sericite alteration. TG3 area. C. Rhyolitic, unsorted, clast‐supported, volcaniclastic rock with pebble‐size clasts and massive sulphide fragments (near the TG3 deposit). D. Green‐grey, feldspar porphyritic dacite with large flow‐foliated chlorite‐quartz ‘pipe’ amygdules. Hanging wall to TG3 deposit. Chapter 2 Page 67 Figure 2.9. Field photographs of mafic rocks from the Cerro El Ereo Formation: A. Typical porphyritic textures of the Cerro El Ereo Formation porphyritic basalt. Sample contains ~20% feldspar phenocrysts to >1 cm in an aphanitic matrix. Non‐amygdaloidal. Subvolcanic or thick flow facies. B. Bleached‐looking, boulder size, subround clasts (C) of basalt feldspar porphyry in a fine matrix of dark grey feldspar porphyritic material (M). Clasts show in‐situ breccia textures (jigsaw‐fit) attributed to progressive fragmentation of blocks during transport. C. Unsorted, non‐stratified basaltic cobble‐ to pebble‐sized lithic and feldspar crystal‐bearing volcaniclastic rock. The sample contains an equal proportion of aphyric to weakly feldspar porphyritic (W) clasts and coarse feldspar (C) porphyry clasts. Amygdaloidal clasts (A) are present but are generally not common. Clast margins often are not easily discernible. D. Thin‐ to thick‐bedded feldspar crystal to ash‐sized tuff; reworked facies at top of formation. Chapter 2 Page 68 Figure 2.10. Field photographs of mafic rocks from the La Bocana Formation: A. Moderately west‐dipping, thick massive basaltic‐andesite flows. Felsic stocks and dykes cut perpendicular to bedding. B. Basaltic andesite dykes with strongly flow foliation defined by flattened and large vesicules (silica amygdules up to 30 cm). C. Polylithic, basaltic‐andesite dominated, mass flow deposit. Not the fractures and in‐situ fragmentation of the clasts due to mass transport (see arrows). D. Cracked and brecciated outer crust of basaltic andesite lava flow and interstitial hyaloclastite resulting from quenching of exposed lava. E. Mafic lobes (M) injected into felsic quartz‐crystal tuffs (T). Tuffs show ‘soft‐sediment’ deformation textures and mafic flows show columnar jointing indicating tuffs were non‐welded/non‐lithified during deposition of mafic flows. F. Lithic and quartz‐feldspar crystal rhyolitic tuff. Colouration of the domains are a result of secondary recrystallization to quartzo‐feldspathic (light) and chloritic (+clay) assemblages due to devitrification of glass component. Chapter 2 Page 69 Figure 2.10 (continued). G. Rhyolitic quartz crystal‐rich and lithic tuff. H. Coarse, boulder breccia with chaotic, unsorted, subround (pillow?) clasts. Talus breccia. I. Medium bedded, well sorted and locally cross bedded (arrow), pebble‐ to sand‐sized, mafic‐dominated volcaniclastic rocks. Chapter 2 Page 70 Figure 2.11. A. thin bedded arenaceous sequence from the Lancones Formation; massive unit at top of outcrop is a diorite sill. B. thin bedded limestones and limey‐arenites. Chapter 2 Page 71 Figure 2.12. Schematic stratigraphic column of the eastern portion of the Lancones basin. Legend as per Figure 2.6. Inset section shows the Tambogrande VMS section in more detail. Age data from this study are shown in their relative stratigraphic positions. Ages of plutonic rocks provided herein (Appendices A, B). Chapter 2 Page 72 Figure 2.13. (following page) 238U/206Pb versus 207Pb/206Pb Tera‐Wasserburg plots (Tera and Wasserburg, 1972) for various volcanic rock samples from the Cerro San Lorenzo and La Bocana Formations. Error ellipses are 2σ. Dashed lines indicate data points omitted versus solid lines/grey ellipses for data included in the age calculation. Inset figures show box plots for all sample points for 207Pb‐corrected 206Pb*/238U data with error bars at 2σ. Open boxes are omitted whereas solid boxes were included in the age calculation. Ages given are 206Pb/238U with 2σ uncertainties. Chapter 2 Page 73 Chapter 2 Page 74 Figure 2.14. 207Pb/235U versus 206Pb/238U U‐Pb concordia plots for various volcanic rock samples from the Cerro San Lorenzo and La Bocana Formations. Chapter 2 Page 75 Figure 2.15. Schematic stratigraphy and U‐Pb zircon ages that constrain the volcanic formations in the Lancones basin. Chapter 2 Page 76 Figure 2.16. Comparison of schematic volcanic stratigraphy of the Lancones Basin and Western Peruvian Trough (modified from Myers, 1974; Offler et al., 1980; Cobbing et al., 1981) with emphasis on age correlation. Legend as per Figure 2.6. Chapter 2 Page 77 2.11 References Alencastre, A.O. 1980. Evaluacion hidrocarburifera de la Cuenca Sechura. Boletin Sociedad Geológica del Perú, 76: 133‐152. Arculus, R.J., Lapierre, H., and Jaillard, E. 1999. 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Batiza R. and White J.D.L. 2000. Submarine lava and hyaloclastite. In Encyclopedia of Volcanoes. Edited by H. Sigurdsson, B. Houghton, S. McNutt, H. Rymer, and J. Stix. Academic Press, New York, pp. 361‐382. Benavides‐Cáceres, V. 1999. Orogenic evolution of the Peruvian Andes: the Andean cycle. In Geology and ore deposits of the central Andes. Edited by B.J. Skinner. Society of Economic Geologists, Special Publication Number 7, pp. 61‐107. Bosch, D., Gabriele, P., Lapierre, H., Malfere, J., and Jaillard, E. 2002. Geodynamic significance of the Raspas metamorphic complex (SW Ecuador); geochemical and isotopic constraints. Tectonophysics, 345: 83‐102. Cas, R.A.F. 1992. Submarine volcanism: eruption styles, products, and relevance to understanding the host rock successions to volcanic‐hosted massive sulfide deposits, Economic Geology, 87: 511‐541. Chávez, A., and Nuñez del Prado, S.H. 1991. Evolución vertical de facies de la serie turbiditica Cretacea (Grupo Copa Sombrero) en el perfíl tipo Huasimal – Encuentros (Cuenca Lancones en el Noreste del Perú): Boletin de la Sociedad Geológica del Perú, 82: 5‐21. Chapter 2 Page 78 Chiaradia, M., Fontbote, L., and Paladines, A. 2004. Metal sources in mineral deposits and crustal rocks of Ecuador (1 degrees N‐4 degrees S); a lead isotope synthesis, Economic Geology and the Bulletin of the Society of Economic Geologists, 99: 1085‐1106. Cobbing, E.J., Pitcher, W.S., Wilson, J.J., Baldock, J.W., Taylor, W.P., McCourt, W., and Snelling, N.J. 1981. The geology of the Western Cordillera of Northern Perú. Institute of Geological Sciences, Natural Environment Research Council, Overseas Memoir 5, 143 p. Daziel, I.W.D. 1981. Back‐arc extension in the southern Andes: A review and critical reappraisal: Philosophical Transactions of the Royal Society of London, Series A, 300: 319‐
335. Franklin, J.M., Gibson, H.L., Jonasson, I.R., and Galley, A.G. 2005. Volcanogenic massive sulfide deposits. In Economic Geology; one hundredth anniversary volume, 1905‐2005. Edited by J.W. Hedenquist, J.F.H. Thompson, R.J. Goldfarb and J.P. Richards. Society of Economic Geologists, pp. 523‐560. Hamilton, W. 1995. Subduction systems and magmatism. In Volcanism Associated with Extension at Consuming Plate Margins. Edited by J. L. Smellie. Geological Society, London, Special Publications, 81, pp. 3–28. Hanson, R.E. and Wilson, T.J. 1991. Submarine rhyolitic volcanism in a Jurassic proto‐marginal basin; southern Andes, Chile and Argentina. Geological Society of America Bulletin, Special Paper 265, pp. 13‐27. Hughes, R.A. and Pilatasig, L.F. 2002. Cretaceous and Tertiary terrane accretion in the Cordillera Occidental of the Andes of Ecuador. Tectonophysics, 345: 29‐48. Jaillard, É, Laubacher, G., Bengston, P., Dhondt, A., and Bulot, L. 1999. Stratigraphy and evolution of the Cretaceous forearc “Celica‐Lancones Basin” of Southwestern Ecuador. Journal of South American Earth Sciences, 12: 51‐68. Jaillard, É, Ordoñez, M., Bengston, P., Berrones, G., Bonhomme, M., Jiménez, N., and Zambrano, I. 1996. Sedimentary and tectonic evolution of the arc zone of Southwestern Ecuador during the Late Cretaceous and Early Tertiary times. Journal of South American Earth Sciences, 9: 131‐40. Jaillard, E. and Soler, P. 1996. Cretaceous to early Paleogene tectonic evolution of the northern central Andes (0‐18 degrees S) and its relations to geodynamics. Tectonophysics, 259: 41‐
53. Jaillard, E., Herail, G., Monfret, T., Diaz‐Martinez, E., Baby, P., Lavenu, A. and Dumont, J.F. 2000. Tectonic evolution of the Andes of Ecuador, Perú, Bolivia and northernmost Chile; tectonic evolution of South America. In Tectonic evolution of South America, 31st International Geological Congress, Rio de Janeiro, Brazil. Edited by U.G. Cordani, E.J. Milani, A. Thomaz Filho and D.A. Campos, pp. 481–559. Chapter 2 Page 79 Jolivet, L., and Tamaki, K. 1992. Neogene kinematics in the Japan Sea region and volcanic activity of the NE‐Japan arc. In Proceedings of the Ocean Drilling Program, Scientific Results, vol. 127/128. Edited by K. Tamaki and K. Suyehiro, J. Allan and M. McWilliams et al. pp. 1311‐1331. Jones, P.R. 1981. Crustal structures of the Perú continental margin and adjacent Nazca Plate, 9 degrees S latitude. Geological Society of America, Memoir 154, p. 423–443. Karig, D.E., Anderson, R.N., and Bibee, L.D. 1978. Characteristics of back arc spreading in the Mariana trough. Journal of Geophysical Research, 83: 1213‐1226. Litherland, M., Aspden, J.A., Jemielita, R.A. 1994. The Metamorphic Belts of Ecuador, Overseas Memoir 11, British Geological Survey, Keyworth, Nottingham, 147 p. Macfarlane, A.W. 1999. Isotopic studies of the northern Andean crustal evolution and ore metal sources. In Geology and ore deposits of the central Andes. Edited by B.J. Skinner, Society of Economic Geologists Special Publication No. 7, pp. 195‐217. Mégard F. 1984. The Andean orogenic period and its major structures in central and northern Perú. Journal of the Geological Society, 141: 893‐900. Mégard, F. 1987. Cordilleran Andes and marginal Andes; a review of Andean geology north of the Arica Elbow (18 degrees S). In Circum‐Pacific orogenic belts and evolution of the Pacific Ocean basin. Edited by J.W.H. Monger and J. Francheteau. American Geophysical Union, Geodynamics Series, Volume 18, p.71‐95. Mitouard, P., Kissel, C., and Laj, C. 1990. Post‐Oligocene rotations in southern Ecuador and northern Perú and the formation of the Huancabamba deflection in the Andean Cordillera. Earth and Planetary Science Letters, 98: 329‐339. Morris, R.C. and Aleman, A.R. 1975. Sedimentation and tectonics of the Middle Cretaceous Copa Sombrero Formation in Northwest Perú. Boletin de la Sociedad Geologica del Perú, 48: 49‐64. Mourier, T., Laj, C., Mégard, F., Roperch, P., Mitouard, P., and Farfau Medrano, A. 1988. An accreted continental terrane in northwestern Perú. Earth and Planetary Science Letters, 88: 182‐192. Myers, J.S. 1974. Cretaceous stratigraphy and structure, Western Andes of Perú between latitudes 10o‐10o30′. The American Association of Petroleum Geologists Bulletin, 58: 474‐
487. Noble, S.R., Aspden, J.A., and Jemielita, R.A. 1997. Northern Andean crustal evolution; new U‐
Pb geochronological constraints from Ecuador. Geological Society of America Bulletin, 109: 789‐798. Chapter 2 Page 80 Offler, R., Aguirre, L., Levi, B., and Child, S. 1980. Burial metamorphism in rocks of the western Andes of Peru. Lithos, 13: 31‐42. Polliand, M., Schaltegger, U., Frank, M., and Fontbote, L. 2005. Formation of intra‐arc volcanosedimentary basins in the western flank of the central Peruvian Andes during late Cretaceous oblique subduction; field evidence and constraints from U‐Pb ages and Hf isotopes. International Journal of Earth Sciences, 94: 231‐242. Reyes, L.R. and Caldas, J.Y. 1987. Geologia de los Cuadranglos de las Playas, La Tina, Las Lomas, Ayabaca, San Antonio. Instituto Geologico Minero y Metalurgio, Bul. 49., 83 p. Scotese, C.R. 1991. Jurassic and Cretaceous plate tectonic reconstructions. Palaeogeography, Palaeoclimatology, Palaeoecology, 87: 493‐501. Sibuet, J., Hay, W.W., Prunier, A., Montadert, L., Hinz, K., Fritsch, J., Hay, W.W., Sibuet, J., Barron, E.J., Brassell, S.C., Dean, W.E., Huc, A.Y., Keating, B.H., McNulty, C.L., Meyers, P.A., Nohara, M., Schallreuter, R.E.L., Steinmetz, J.C., Stow, D.A.V., Stradner, H., and Boyce, R.E. 1984. Early evolution of the South Atlantic Ocean; role of the rifting episode; initial reports of the deep sea drilling project covering leg 75 of the cruises of the drilling vessel Glomar Challenger, Walvis Bay, South Africa to Recife, Brazil, July‐September, 1980, Initial Reports of the Deep Sea Drilling Project, 75: 469‐481. Soler, P. and Bonhomme, M.G. 1990. Relation of magmatic activity to plate dynamics in central Peru from Late Cretaceous to present. In Plutonism from Antarctica to Alaska. Edited by S.M. Kay and C.W. Rapela. Geological Society of America Special Paper 241, pp. 173‐192. Spikings, R.A., Winkler, W., Hughes, R.A. and Handler, R. 2005. Thermochronology of allochthonous terranes in Ecuador: Unraveling the accretionary and post‐accretionary history of the Northern Andes. Tectonophysics, 399: 195‐220. Steinmüller, K., Chacón Abad, N., and Grant, B. 2000. Volcanogenic massive sulphide deposits in Perú. In VMS Deposits of Latin America. Edited by R. Sherlock and M.A.V. Logan. Geological Association of Canada, Mineral Deposits Division, Special Paper No.2. pp. 423‐
437. Taylor, B.J., Goodliffe, A., Martinez, F., and Het, R. 1995. Continental rifting and initial seafloor spreading in the Woodlark Basin. Nature, 374: 534‐537. Tegart, P., Allen, G., Carstensen, A. 2000. Regional setting, stratigraphy, alteration and mineralization of the Tambo Grande VMS district, Piura Department, Northern Perú. In VMS Deposits of Latin America. Edited by R. Sherlock and M.A.V. Logan. Geological Association of Canada, Mineral Deposits Division, Special Paper No.2. pp. 375‐405. Tera, F., and Wasserburg, G.J. 1972. U‐Th‐Pb systematics in three Apollo 14 basalts and the problem of initial Pb in lunar rocks. Earth and Planetary Science Letters, 14: 281‐304. Chapter 2 Page 81 Winter, L.S., Tosdal, R., Franklin, J.M., Tegart, P. 2004. A Reconstructed Cretaceous Depositional Setting for Giant Volcanogenic Massive Sulfide Deposits at Tambogrande, Northwestern Perú. In Special Publication 11 ‐ Andean Metallogeny: New Discoveries, Concepts, and Updates, Edited by R.H. Sillitoe, J.Perello, and C.Vidal. Society of Economic Geologists, pp. 319‐340. Winter, L.S., Tosdal, R.M., Tegart, P. 2002. A step in the formation of the Huancabamba deflection in the Andes of Perú and Ecuador. Geological Society of America Annual Meeting. Paper No. 191‐14. White, J.D.L., Smellie, J.L., and Clague, D.A. 2003. Introduction: A deductive outline and topical overview of subaqueous explosive volcanism. In Explosive subaqueous volcanism. Edited by J.D.L. White, J.L. Smellie, and D.A. Clague. American Geophysical Union Geophysical Monograph 140, pp. 1–23. Vergara, M., Levi, B., Nyström, J.O., Cancino, A. 1995. Jurassic and Early Cretaceous island arc volcanism, extension, and subsidence in the Coast Range of central Chile. Geological Society of America Bulletin, 107: 1427‐1440. Chapter 2 Page 82 Chapter 3. A Reconstructed Cretaceous Depositional Setting for Giant VMS Deposits at Tambogrande, NW Perú 2 3.1 Manuscript Status A version of this chapter has been published. Only the abstract is provided here due to Copyright limitations. 3.2 Abstract The Cretaceous Tambogrande volcanogenic massive sulphide (VMS) deposits of northern Perú are amongst the largest Cu‐Zn‐Au‐Ag bimodal‐mafic VMS deposits in the world. There are three known deposits each with approximately 100 million metric tons (Mt) of pyrite‐rich massive sulphide. The deposits are intimately associated with dacite lava dome complexes and were deposited within steep‐sided basins on the seafloor. Reconstructed seafloor paleomorphological models indicate sulphide deposition was focussed in the deepest parts of the basins. Sulphide deposition accompanied syn‐volcanic faulting and episodic dacitic and basaltic eruptions. A series of time‐stratigraphic horizons are noted at TG1 and TG3 and mark stages in the development of the volcanic complex and massive sulphides. There is only limited evidence for replacement of host rocks during formation of the Tambogrande deposits, unlike many other large massive sulphide deposits. The deposits at Tambogrande resulted from focused hydrothermal fluid flow along syn‐volcanic faults and the deposition of sulphides within relatively deep and restricted basins. These depressions of up to several hundred metres deep, the result of the structural and volcanological setting, acted as efficient ‘traps’ for sulphide 2
A version of this chapter has been published. Winter, L.S., Tosdal, R., Franklin, J.M., Tegart, P. 2004. A
Reconstructed Cretaceous Depositional Setting for Giant Volcanogenic Massive Sulfide Deposits at Tambogrande,
Northwestern Perú. In Special Publication 11 - Andean Metallogeny: New Discoveries, Concepts, and Updates.
Edited by R.H. Sillitoe, J.Perello, and C.Vidal. Society of Economic Geologists, pp. 319-340.
Chapter 3 Page 83 deposition and also were also important for the preservation of the sulphide, as they acted to shield the accumulations from potential submarine oxidation and weathering. Steep basins and episodic bimodal eruptions are key geologic attributes of the depositional setting at Tambogrande and may be necessary for the formation of anomalously large VMS deposits in a volcanic rock‐dominated environment. Chapter 3 Page 84 Chapter 4. Volcanic Rock Geochemistry and the Geodynamic Setting of VMS Deposits at Tambogrande, Perú 3 4.1 Overview ‘Giant’ volcanogenic massive sulphide (VMS) deposits in the Tambogrande area of northwestern Perú occur along a single time‐stratigraphic horizon within the ca 105 Ma Cerro San Lorenzo Formation, the lowermost volcanic sequence deposited within the Lancones arc‐
rift basin. The Lancones basin is part of a larger series of segmented, continental margin rift basins extending from southwestern Ecuador and along the Peruvian coast. The three known deposits (TG1: 109 Mt grading 1.6% Cu, 1.0% Zn, 0.5 g/t Au and 22 g/t Ag, plus 16.7 Mt grading 3.5 g/t Au and 64 g/t Ag in oxide ore; TG3: 82 Mt grading 1.0% Cu, 1.4% Zn, 0.8 g/t Au and 25 g/t Ag, and B5: resource not defined) constitute the most significant VMS camp in South America. A lithogeochemical study of the volcanic arc sequence of the Lancones basin illustrates that all volcanic units sampled are ‘arc‐related’ and display prominent negative Nb anomalies relative to Th and Nb on primitive‐mantle normalized trace element plots. The VMS‐hosting Cerro San Lorenzo Formation is a bimodal sequence dominated by basaltic lava flows of transitional to mildly calc‐alkaline affinity defined by moderate REE abundances (YbN = 4–12), moderate LREE/HREE fractionation [(La/Yb)N = 0.6–7.3], low‐moderate Zr/Y ratios (1.5‐5.3), and negative HFSE (high field strength element) anomalies (e.g., Nb/Nb* = 0.1–0.5; Zr and Hf display slight depletions). Felsic volcanic rocks in this formation are M‐type (i.e., mantle‐derived; low Nb, Y) and are comparable in geochemical affinity to the basalt regarding LREE‐profiles [(La/Yb)N = 1.2‐6.6], Zr/Y ratios (3.0‐7.1), and negative Nb anomalies (Nb/Nb* = 0.1–0.3). However, the felsic rocks differ in having higher REE abundances (YbN = 10‐24), some HFSE 3
A version of this chapter will be submitted for publication. Winter, L.S. and Tosdal, R., Volcanic Rock
Geochemistry and the Geodynamic Setting of VMS Deposits at Tambogrande, Perú.
Chapter 4 85 enrichment (Zr/Zr* = 1.0–2.2; Hf/Hf* = 1.1–2.5), and pronounced negative Ti (Ti/Ti*= 0.1‐0.2) and Eu anomalies (Eu/Eu* = 0.4–0.9). The rhyolites range from FII to FIII‐type. Overlying the Cerro San Lorenzo Formation, the Cerro El Ereo Formation comprises a uniform sequence of coarse feldspar porphyritic basalt flows and breccia with transitional affinities. These are characterized by negligible LREE/HREE fractionation [(La/Yb)N = 0.4‐1.6] and low Zr/Y ratios (1.4 – 2.8). The uppermost volcanic sequence, the La Bocana Formation, is a bimodal sequence characterized by transitional to tholeiitic mafic volcanic rocks [(La/Yb)N = 0.5‐
2.6, Zr/Y = 1.3‐3.3]. M‐type rhyolites [(La/Yb)N = 0.5‐2.6, Zr/Y = 1.3‐3.3] are very similar to the Cerro San Lorenzo Formation but are mostly FIV‐like. There are only subtle variations in the petrochemistry of the arc volcanic sequence in the Lancones basin which favour VMS formation in the early rift, Cerro San Lorenzo Formation, which include: (i) Evolution from mildly calc‐alkaline to tholeiitic affinities; this possibly reflects a depletion of the mantle source with time. A relatively static mantle wedge with limited replenishment of fertile asthensophere during waning of the extensional phase of the rift event may be responsible. Moreover, the absence of MORB‐type lavas suggests the rift basin did not evolve to back‐arc rifting. VMS formation was more likely in the early stages of arc volcanism during the initial rifting and period of maximum extension; (ii) Average basalt composition is ‘less fractionated’ in the Cerro San Lorenzo Formation, with higher MgO, Ni, Cr, than basalt in the overlying Cerro El Ereo and La Bocana formations. This may be related to the presence of thinner ocean crust in the basin during the early stages of rifting and the concomitant more rapid transfer of basaltic magma to surface. This reflects the greater heat transfer to shallow crustal levels in the early stages of arc volcanism, is consistent with the timing of VMS formation in the Lancones basin; and (iii) M‐type felsic volcanic rocks have lower HFSE and REE Chapter 4 86 contents in the La Bocana Formation than in the Cerro San Lorenzo Formation. In the immediate vicinity of the VMS deposits, rhyolite does not define fractionation trends with the mafic rocks, rather, the variable REE and HFSE ratios indicate contamination of the mafic‐
derived lavas with an incompatible element enriched source (i.e., continental crust). These felsic magmas were likely partial melts of upper crustal portions of the juvenile arc, or were primary melts of the mantle‐wedge that assimilated older crust. The ore‐associated felsic volcanic rocks typify the generally more prospective FII to FIII‐type rhyolites, though with slightly lower average HFSE and REE contents. Rhyolite in the VMS‐barren La Bocana Formation is dominantly FIV type. This suggests the petrogenesis of felsic volcanic rocks associated with Phanerozoic VMS deposits is not substantially different than that of some Archean VMS systems. The ocean arc‐type geochemical affinities which define the Lancones basin volcanic rocks are unusual considering the continental margin geodynamic setting, where continental rift or continental back‐arc rift affinities might be expected. Since the arc does not appear to have had a protracted history of extension and did not fully evolve to a back‐arc, the pre‐Andean (Early Cretaceous) continental crust must have been relatively thin prior to arc magmatism. Mantle‐derived magmas traversed the attenuated crust and preserved their relatively immature island arc geochemical characteristics. 4.2 Introduction Detailed volcanic rock geochemical studies in ancient volcanic settings are used to understand the petrogenesis of a volcanic suite and in turn can determine the relative prospectivity to host volcanogenic massive sulphide (VMS) mineralization (Lesher et al., 1986; Swinden et al., 1997; Lentz, 1998; Bailes and Galley, 1999; Barrett and MacLean, 1999; Piercey Chapter 4 87 et al., 2001). However, studies of the tectonic history of VMS‐hosting terranes are hindered in some ancient successions by tectonism, which has complicated elucidation of geologic events. Nonetheless, in many VMS districts there is strong petrochemical evidence to support a spatial relationship between extensional arc geodynamics and VMS formation (e.g., Swinden, 1991; Lentz, 1998; Piercey, 2007). This link is based largely on volcanic rock lithogeochemistry and anologies to recent volcanic rock suites in modern settings. The Cretaceous Lancones basin is a relatively young volcanic terrane compared with most other VMS‐bearing terrains, and hence its geodynamic framework is relatively well understood. Plate tectonic models for the South American continent during the Cretaceous (e.g., Scotese, 1991) and tectonic evolution of the Andes in particular (Benavides‐Cáceres, 1999; Jaillard et al., 2000) are relatively well constrained. Stratigraphic and geochronological control on the volcanic successions of the Lancones Basin is also well established (Jaillard et al., 1999; Chap. 2). At Tambogrande, arc‐related rifting and volcanism at the continental margin is associated with VMS formation. This chapter provides a baseline whole rock lithogeochemical survey for volcanic rocks of the Lancones basin and a detailed chemostratigraphic model for volcanic lithologies associated with VMS deposits at Tambogrande. The Lancones basin is located in the coastal regional of northwestern Perú and extends into southwestern Ecuador. The basin represents one of several marginal rift basins along the Peruvian continental margin that are host to numerous VMS deposits (Fig. 4.1). However, the Lancones basin hosts the most significant of all VMS deposits in the Andes at Tambogrande (Fig. 4.2). The Tambogrande district (Injoque et al., 1979; Tegart et al., 2000; Chapter 2) is defined by three anomalously large (~100 Mt) base‐ and precious‐metal‐bearing VMS deposits (Chap. 1). The deposits are within the upper 3% of all VMS deposits of their type globally in terms of Chapter 4 88 total metal content (Franklin et al., 2005) and can be classified as bimodal mafic type according to Barrie and Hannington (1999). Analogous VMS deposits are of Archean age (e.g., Noranda, Gibson and Watkinson, 1990; Kidd Creek, Hannington and Barrie, 1999), Proterozoic age (e.g., Flin Flon, Syme and Bailes, 1993) and Phanerozoic age (e.g., Sibai, Gai; Herrington et al., 2005). Cretaceous marginal basins along the western margin of South America have been ascribed to various tectonomagmatic scenarios ranging from ‘aborted ophiolite’ (Aguirre and Offler, 1985), Gulf of California‐type continental rifts (Atherton, 1990), modified back‐arc spreading centres, with the potential arc lying west of the present coast (Atherton et al, 1983; Petford and Atherton, 1995), arc models associated with a steeply dipping subduction zone (Soler and Bonhomme, 1990), and accreted island arc or back‐arc (Wise, 2000). Whereas the Western Peruvian Trough of central Perú is partly eroded along its western margin, the Lancones basin in northwestern region of Perú and southwestern Ecuador includes both the volcanic arc and the forearc to the west. The allochthonous Amotape complex (Mourier et al., 1988; Noble et al., 1997) and the Mesozoic continental margin Olmos complex (Reyes and Caldas, 1987) are preserved at the present day margin of the volcanic arc sequences. With all of the tectonostratigraphic components preserved, the Lancones basin provides the best opportunity to study the tectonic history and magmatism of the Cretaceous marginal basins of Perú‐
Ecuador. In this chapter the petrogenesis of the volcanic rocks and the relationship to VMS formation is investigated. 4.3 Tectonic Setting The Tambogrande VMS deposits formed in the Lancones Basin in association with the deposition of mid‐Cretaceous volcano‐sedimentary strata (Reyes and Caldas, 1987; Chap. 2.). The Lancones Basin represents the northern portion of a much larger and segmented Mesozoic Chapter 4 89 continental margin rift basin extends from southwestern Ecuador (Jaillard et al., 1996) along the Peruvian coast (Western Peruvian Trough or Huarmey‐Cañete Marginal Trough; Atherton et al., 1983; Benavides‐Cáceres, 1999). Similar arc‐related marginal rifts occur in Chile (Coast Range; Vergara et al., 1995) and Argentina (Rocas Verdes basin; Dalziel, 1981; Hanson and Wilson, 1991). These marginal basins represent an episode of extension in the first phase (~Mesozoic) of the Andean cycle in which Mariana‐type subduction permitted crustal attenuation, deposition of major marine sequences and eruption of large volumes of mafic‐
dominated, subduction‐related volcanic rocks (Benavides‐Cáceres, 1999). The rift regime is temporally linked to the opening of the South Atlantic in the Cretaceous and lasted until the Late Cretaceous when the geodynamical cycle shifted towards Andean‐type subduction in part due to active spreading in the South Atlantic (Soler and Bonhomme, 1990). This resulted in the termination of marine sedimentation and the beginning of contractional tectonism, continental arc magmatism and a tectonic regime which continues to the present day (Benavides‐Cáceres, 1999; Jaillard et al., 2000) . The Lancones basin is located within the Huancabamba deflection, a major oroclinal bend in the Andes which separates the north‐northwest‐trending Peruvian Andes from the northeast‐
trending Ecuadorian Andes (Fig. 4.1). During the Mesozoic this region records a sequence of accretionary events, shifts in convergence direction, and contractional and extensional rotations (Mitouard et al., 1990). In the Jurassic, a southeast‐directed subduction zone was responsible for continental arc volcanism along the Ecuadorian segment (Litherland et al., 1994), whereas a sinistral transform system occurred along the Peruvian segment (Jaillard et al., 2000). During the Early Cretaceous a convergence shift to the northeast terminated the Ecuadorian magmatic arc and established the conditions for subduction along the Peruvian Chapter 4 90 segment (Jaillard et al., 2000). This tectonic development along the Peruvian‐Ecuadorian margin in the Early Cretaceous is synchronous with a rift event and the formation of the Lancones basin. Several accretionary events beginning in the Early Cretaceous also played a major role in the development of the margin, especially in northwestern Perú and Ecuador where allochthonous terranes are identified (Litherland et al., 1994). The Amotape terrane represents a microcontinental block comprised of metamorphosed sedimentary and plutonic rocks with Middle to Late Triassic U‐Pb ages (Noble et al., 1997; Appendix A), as well as high‐
pressure metamorphosed oceanic terranes thought to represent mid‐ocean ridge and oceanic plateau basalt (e.g., Raspas metamorphic complex; Arculus et al., 1999; Bosch et al., 2002). Mourier et al. (1988) suggest the Amotape terrane arrived from the south, the timing of which is constrained by cooling ages of ~132‐110 Ma (Bosch et al., 2002). The accretionary event likely triggered the initiation of the new subduction zone outboard of the Amotape terrane resulting in the formation of the Lancones basin. Subsequent rifting of this basin probably utilized the suture between the Amotape Terrane and continental South America (Litherland et al., 1994) . 4.4 Regional Geology and Volcanic Stratigraphy The Lancones Basin is limited to the east, west and north by continental crustal blocks that represent the Jurassic to Early Cretaceous pre‐rift Andean margin (Benavides‐Cáceres, 1999). Geophysical models constrain the crustal architecture of the continental margin of Perú and demonstrate a large arch‐like structure of dense (3.25 g/cm3) material that coincides with the Mesozoic volcanic rift sequences, interpreted to be oceanic crust and separating continental blocks (Couch et al., 1981; Jones, 1981). These continental blocks were topographic highs during marine deposition in Mesozoic times (Cobbing et al., 1981). To the southeast of the Chapter 4 91 Lancones Basin, the Paleozoic(?) Olmos Massif is a probable reactivated margin of the Amazonian craton (Macfarlane, 1999). This poorly characterized terrane consists of pre‐
Ordovician greenschist facies pelitic to psammitic rocks overlain by platform carbonate rocks of Triassic to Early Jurassic age. The Olmos complex is considered equivalent to Marañon Geanticline (Cobbing et al., 1981; Reyes and Caldas, 1987; Mourier et al., 1988; Litherland et al., 1994). Bordering the Lancones basin to the southwest, northwest and north is the Amotope terrane, comprised of Mesozoic gneisses, meta‐sedimentary and meta‐volcanic oceanic rocks (Mourier et al., 1988; Aspden et al., 1995; Litherland et al., 1994; Arculus et al., 1999). Rocks of the Lancones Basin are exposed in northwestern Perú and southwestern Ecuador for more than 135 kilometres along a northeast trend and approximately 150 kilometres across the trend. Tertiary cover blankets the basin in the southwest for an additional 50 kilometres. The basin can be subdivided in an eastern volcanic arc and western sedimentary forearc. The volcanic arc sequence, up to 80 km wide, is dominated by submarine mafic volcanic and volcaniclastic rocks, and grades into forearc sedimentary rocks which dominate the western portion of the Lancones Basin (Jaillard et al., 1999). A ~3 km thick turbidite sequence, the Copa Sombrero Group (Chávez and Nuñez del Prado, 1991; Morris and Aleman, 1975; Jaillard et al, 1996, 1999), represents the forearc and temporally overlaps the volcanic arc sequence. The four formations of the volcanic arc sequence in the Lancones basin have a cumulative thickness of ~ 8 to 10 km and are subdivided into two main tectono‐volcanic phases based on depositional facies, composition and chronology (Figs. 4.4, 4.5; Chap. 2). The Cerro San Lorenzo Formation represents phase 1 and the Cerro El Ereo, La Bocana and Lancones formations record phase 2. Phase 1 is a mafic‐dominated sequence characterized by lava flows and associated breccias, minor aphyric felsic lavas, and a general absence of siliciclastic sedimentary rocks. The Chapter 4 92 rocks are interpreted to have been deposited in a relatively deep water environment. All VMS deposits in the Lancones basin are hosted by the phase 1 volcanic sequence. The phase 2 volcanic cycle is an 8 km‐thick sequence of mafic to felsic volcanic and volcaniclastic rocks with subordinate calcareous and siliciclastic sedimentary rocks. These rocks represent in general a relatively more shallow water setting. The upper part of the sequence includes re‐worked volcaniclastic rocks which grade into sedimentary rocks in the upper sectors marking a transition to forearc turbidites of the Copa Sombrero Group. In Ecuador the volcanic arc sequence has not been studied in detail and is described as a 2 to 3 km‐thick package of dominantly mafic pillow lavas and related volcaniclastic rocks (Jaillard et al., 1996). Metamorphic grades range from zeolite to lower greenschist assemblages, with the higher metamorphic grades occurring near plutonic rocks. Diagenesis and low‐temperature, near seafloor metasomatism, due to ambient seawater‐rock interaction (halmyrolosis), have resulted in a wide range of low‐temperature replacement and open‐space‐filling minerals within basaltic rocks, including analcite, albite (after plagioclase), amphibole (uralitized clinopyroxene), carbonate minerals, chlorite, epidote, hematite, palagonite (after groundmass glass), prehnite, pumpellyite, sericite and/or various clays (sausseritized feldspar), and zeolites. Hydrothermal alteration is confined to discordant zones of the footwall rocks immediately below massive sulphide deposits and includes variable replacement of the rocks by chlorite, sericite and quartz in addition to stringer and disseminated sulphide mineralization (Tegart et al, 2000). None of the rocks in this study show the effects of dynamic metamorphism and primary textures are generally well preserved. 4.4.1 Cerro San Lorenzo Formation The lowermost volcanic sequence, the Cerro San Lorenzo Formation, is characterized by bimodal volcanic rocks and is dominated by basalt. The depositional environment is inferred to Chapter 4 93 have been a relatively deep marine based on the absence of pyroclastic rocks and the presence of minor black mudstone as the sedimentary rocks. Felsic volcanic rocks represent a minor component of the total estimated 2,500 metre thickness of the formation. Volcanic rocks of the Cerro San Lorenzo Formation host all VMS deposits and prospects in the Tambogrande district (TG1, TG3, and B5). U‐Pb zircon ages of 99.8 to 104.8 Ma for syn‐mineralization rhyolitic volcaniclastic rocks and post‐mineralization rhyolite dykes, respectively, have been determined (Chap. 2). Basalt in this formation is typically feldspar‐augite porphyritic, amygdaloidal and occurred as massive to pillowed flows associated with a variety of autoclastic breccias, ranging from pillow fragment and talus breccia, hyaloclastite breccias and globular autobreccia. Felsic volcanic rocks are represented by massive to flow‐banded lavas, domes or dykes of dacitic to rhyolitic composition and are associated with autoclastic proximal volcaniclastic rocks. A rhyolite flow‐dome complex, up to 2 km in extent by 250 m in composite thickness, is dominated by buff to light grey, aphyric to weakly feldspar porphyritic, locally spherulitic and commonly perlitic rhyolite. Plagioclase and resorbed quartz porphyritic rhyolite dykes cut massive sulphide deposits (e.g., TG1). Quartz porphyritic rhyolite lavas, however, are not recognized within the Cerro San Lorenzo Formation. A few amygdule‐rich dacitic lava flows and breccia deposits are conspicuous in the immediate hanging wall of the TG1 and TG3 deposits. Volumetrically minor pelagic sedimentary rocks are ubiquitous throughout this formation. 4.4.2 Cerro El Ereo Formation The Cerro El Ereo Formation is an entirely mafic volcanic sequence of distinct coarse feldspar porphyritic lava and minor breccia. Lavas are generally massive with few amygdales. Volcaniclastic rocks are typically thick‐bedded, non‐stratified, boulder‐sized breccia, though Chapter 4 94 fairly well sorted lithic and feldspar crystal tuff of mafic composition are present near the upper contact with the La Bocana Formation. 4.4.3 La Bocana Formation The La Bocana Formation is a bimodal basalt/basaltic‐andesite and rhyolitic succession that is notably volcaniclastic‐rich. Mafic rocks in the formation include highly vesicular, thick flows and dykes with well developed flow foliations grading into autoclastic deposits. Felsic rocks range from quartz‐ and/or feldspar‐porphyritic flows to crystal, lithic, and pumice tuffs. The pyroclastic rocks include cross bedded volcanic sandstones which indicate a relatively shallow water depositional setting. U‐Pb zircon ages range from 99.3‐91.1 Ma (Chap. 2). 4.4.4 Lancones Formation The Lancones Formation consists of a polylithic, mafic‐dominated, volcaniclastic units deposited within in a relatively shallow marine environment. Thick‐bedded and variably stratified breccias grade upwards into siliciclastic and calcareous sedimentary units (Chapter 2). Fossils in the age range of late Albian and Early Cenomanian are reported in this formation (Reyes and Caldas, 1987). These rocks mark the transition to forearc sedimentary sequences of the Copa Sombrero Group in the western Lancones Basin, though the contact is defined to lie west of the study area (Jaillard et al., 1999). 4.5 Lithogeochemistry 4.5.1 Sampling Procedures and Analytical Methods One hundred and twenty‐nine whole rock samples were collected from volcanic and sub‐
volcanic units, including outcrop and diamond drill core, Samples from the 2000 field season, mostly in the vicinity of the VMS deposits, were analyzed at Bondar‐Clegg Laboratories in Vancouver, British Columbia, Canada using lithium metaborate fusion, inductively coupled plasma – atomic emission spectrometry (ICP‐AES). Pressed powder pellet X‐ray fluorescence (XRF) and for additional trace and rare earth element analyses were completed by using Na2O2 Chapter 4 95 sinter, inductively coupled plasma – mass spectrometry (ICP‐MS) at Memorial University in St. John’s, Newfoundland, Canada. Samples collected during 2002 regional mapping program from outcrop and drill core were analyzed for major and trace elements by a combination of lithium metaborate fusion ICP‐AES and ICP‐MS and pressed powder pellet XRF at ALS‐Chemex Laboratories in Vancouver. The complete dataset, locations and rock types, as well as details on the analytical technique, precision and accuracy are given in Appendix C. 4.5.2 Alteration and Element Mobility Element mobility as a result of mineralogical changes during fluid‐rock reactions is a feature of nearly all submarine volcanic rocks, either as a result of halmyrolysis or burial diagenesis. Variable mobility of many major elements, principally Ca, K, Na, Si, Mg, and Fe is noted, whereas Al, Ti, and P are generally immobile in VMS‐related altered rocks (Hajash and Chandler, 1981; Saeki and Date, 1983; Barrett and MacLean, 1999; Large et al., 2001) and indeed exhibit conservative behavior in this data set. The trace element suites also display variable mobility, with low field strength elements (LFSE; e.g., Ba, Rb, Sr) considered highly mobile due to a chemical behavior similar to Ca and K. The exception is Th, which is highly immobile. The high field strength elements (HFSE; e.g., Ga, Hf, Nb, Ta, Y, Zr) are considered highly immobile as they are usually hosted in accessory minerals not easily affected by VMS‐related hydrothermal fluids (e.g., titanite, zircon, monazite, xenotime). The rare earth element (REE) group is generally immobile in relatively weakly altered (e.g., sericitized) rocks but can be mobilized under conditions yielding chlorite‐rich alteration assemblages (Campbell et al., 1984). Divalent Eu is the exception within the REE group due to its ability to substitute for alkalies in feldspar. Sc, V and Cr are also generally considered relatively immobile. In this study mobile elements are used on several plots but some scatter due to alteration is noted. Plots utilizing mobile elements are used to make broad characterizations, whereas most of the conclusions of the Chapter 4 96 report are supported on immobile trace element and trace element ratio plots. ‘Least altered’ samples are used in all cases. 4.5.3 Geochemical Results Samples are grouped according to their stratigraphic position within the Cerro San Lorenzo, Cerro El Ereo and La Bocana Formations. No samples were collected from the Lancones Formation for lithogeochemical analyses as it is dominated by volcaniclastic and sedimentary rocks. Samples of felsic volcanic rocks within the Cerro San Lorenzo Formation are further subdivided based on facies and their stratigraphic location relative to VMS deposits (i.e., syn‐
mineralization rhyolite, post‐mineralization dacite, and rhyolite dykes). Most samples represent coherent facies of volcanic rocks, except some samples are larger clasts within coarse volcaniclastic units that are interpreted as relatively proximal deposits. Samples were collected throughout the map area and represent a stratigraphic section up to 8 km‐thick and from a large geographic area (>3,000 km2), though there is a higher sample population density in the Tambogrande area (Fig 4.4). Summarized geochemical results for basalt, basaltic andesite, dacite and rhyolite from each formation are given in table 1. Key trace element ratios are provided in table 2. 4.5.3.1 Cerro San Lorenzo Formation The Cerro San Lorenzo Formation includes mostly basaltic and dacitic to rhyolitic compositions (Fig. 4.6 A, B), with subordinate intermediate composition rocks. Low total Na2O and K2O as well as Nb/Y < 0.5 indicate subalkaline affinities. On the AFM plot, the samples define a trend that mostly approximates the calc‐alkaline field, though the mafic rocks overlap the tholeiitic field as well (Fig 4.6C). Basalt, defined by < 55% SiO2 and Zr/TiO2 < 106, contains between ~5.0 to 11.5 wt.% MgO for most samples, which is comparatively higher than either of the Cerro El Ereo or La Bocana Chapter 4 97 formations (Fig 4.7A). Though moderate TiO2 values are similar within basalt from all formations, basalt from Cerro San Lorenzo Formation yields relatively high Ni and Cr values (Fig 4.7B,C) and moderate to low V values (Fig 4.7D). On various basalt discrimination diagrams these rocks are defined as arc volcanic rocks of broadly calc‐alkaline affinity (Fig 4.8 A‐C), with V/Ti ratios typical of oceanic arc or back‐arc settings (Fig 4.8d). REE abundances span a narrow range (YbN = 4‐13). Light REE (LREE) fractionation is evident in the dominantly negative slope of the normalized REE profiles (Fig 4.9A) and is emphasized by dominantly positive (La/Yb)N 4 values of 0.6 to 7.3 and mimicked by Zr/Y values (1.5 – 5.3). Slightly more evolved rocks, i.e., basaltic‐andesite, have similar normalized trace element profiles similar to basalt but display negative Ti, Al, Sc, and V depletion and Zr and Hf enrichment, as well as strong LREE fractionation (Fig 4.9B). A pronounced negative Nb anomaly characterizes all normalized trace element profiles. Basalt in this formation has HFSE and LFSE abundances characterized by low Nb (Nb/Nb* = 0.06‐0.5), low Ti (Ti/Ti* = 0.5‐1.1), moderate Zr (Zr/Zr* = 0.6‐1.3), and relatively high Th [(Th/La)PM 5 = 1‐4]. (La/Yb)N and Zr/Y values of basalt from the Cerro San Lorenzo Formation are distinctively higher than either the Cerro El Ereo or La Bocana formations (Fig. 4.10A, B). Low HFSE contents for felsic rocks are reflected in Zr versus Ga/Al and Nb versus Y plots (Fig. 4.11 A‐B). On these discrimination plots, felsic rocks are categorized dominantly as M‐type whereas some samples plot closer to the I‐type field (M‐type or direct mantle‐type volcanic arc rocks are generally considered to be derived from a mafic parent rock whereas I‐type are interpreted to be recycled mantle material due to partial melting of juvenile crust; Pitcher, 1983; Whalen, 1985). Primitive mantle‐normalized extended trace element plots (Fig. 4.12A) 4
5
N
denotes chondrite-normalized value of Sun and McDonough (1989).
denotes primitive mantle-normalized value of Sun and McDonough (1989).
PM
Chapter 4 98 show broadly similar patterns for all felsic rocks from the Cerro San Lorenzo Formation. Pronounced negative Ti (Ti/Ti*=0.1 to 0.2) and Nb anomalies (Nb/Nb*=0.1‐0.3) are ubiquitous. Transition metals (e.g., Al, Sc, V) are also strongly depleted. With one exception, this group displays moderate LREE fractionation [(La/Yb)N = 1.2‐6.6)] and moderate Zr/Y ratios (3.0‐7.1), consistent with transitional affinities. All samples display positive Zr (Zr/Zr* = 1.0‐2.2) and Hf (Hf/Hf* = 1.0‐2.5) anomalies. LFSE‐enrichment is evident in the high Th contents relative to LREE [(Th/La)PM = 1.5‐5.7]. Strong Eu depletion is evident in all samples (Eu/Eu* = 0.4‐0.9). 4.5.3.2 Cerro El Ereo Formation The Cerro El Ereo Formation includes mafic compositions only (basalt, minor basaltic‐
andesite) with relatively low alkali contents and low Nb/Y (~0.04), indicating a strongly subalkaline affinity (Fig. 4.6 A,B). All samples plot within the tholeiitic field on the AFM diagram (Fig. 4.6C). Basalt yields relatively low MgO, Ni, and Cr suggesting a relatively fractionated suite (Fig. 4.7A‐C), and moderate‐high V contents are consistent with a tholeiitic affinity (Fig 4.7d). Relatively low HFSE and LFSE contents and high V/Ti ratios in basalt are also consistent with island arc tholeiites on basalt discrimination diagrams (Fig 4.8A‐D). REE abundances are similar to basalt from the Cerro San Lorenzo formation (Yb* = 4‐12) but differ on normalized trace element plots in that relatively flat REE profiles indicate only weak LREE fractionation [(La/Yb)N = 0.4‐1.6] (Fig 4.9C). Basalt in this formation is characterized by relatively low Nb (Nb/Nb*=0.1‐
0.4), low Ti (Ti/Ti* = 0.6‐1.7), moderate Zr (Zr/Zr*=0.6‐1.3), and relatively low Th [(Th/La)PM < 1]. One sample of basaltic‐andesite does not vary significantly from basalt with respect to the trace element profile (Fig 4.9D). 4.5.3.3 La Bocana Formation The La Bocana Formation spans a range of compositions from basalt to rhyolite. These rocks are characterized as subalkaline (low alkali contents, Nb/Y ~ 0.1; Fig 4.6 A‐B), and overlap Chapter 4 99 both the tholeiitic (for most mafic rocks) and calc‐alkaline fields (for most felsic rocks) on the AFM diagram (Fig 4.6C). Basalt from the La Bocana Formation has the lowest average MgO, Ni and Cr (Fig 4.7A‐C) and moderate to high V contents (Fig 4.7D). HFSE and LFSE abundances and V/Ti ratios are most similar to the Cerro El Ereo Formation, overlapping both the calc‐alkaline and tholeiitic fields on basalt discrimination diagrams (Fig 4.8A‐D). On normalized trace element and REE plots basalt yields HFSE‐depletion with prominent negative Nb (Nb/Nb* ~ 0.1), Ti (Ti/Ti* = 0.5‐1.4), and Zr anomalies (Zr/Zr* = 0.6‐1.0) and LFSE‐enrichment represented by Th [(Th/La) PM =1.1‐3.2] (Fig 4.9E). REE abundances are similar to the other formations (YbN = 6 ‐ 13). Most samples display slight depletion to slight enrichment of LREE relative to HREE [(La/Yb)N = 0.5 to 2.6] and relatively low Zr/Y (1.3 – 3.0) in comparison to basalt in the above mentioned formations. Basaltic‐andesite of the La Bocana Formation yields normalized trace element profiles quite different from associated basalt within the same formation. The notable differences include enrichment of Zr and Hf, depletion of Ti, and strong LREE fractionation [(La/Yb)N = 2.0‐8.0] (Fig 4.9F) in the basaltic andesites. These variations are difficult to explain by fractionation of a basaltic magma alone and suggest crustal contamination. Felsic volcanic rocks from the La Bocana Formation are broadly geochemically similar to those of the Cerro San Lorenzo Formation. These rocks yield low HFSE contents and in Zr versus Ga/Al and Nb versus Y plots fall within the field of M‐type volcanic arc rocks (Fig 4.11 A, B). Primitive mantle‐normalized extended trace element plots (Fig 4.12B) are fairly uniform for all felsic volcanic rocks from this formation and display pronounced negative Ti (Ti/Ti* = 0.04 to 0.3) and Nb anomalies (Nb/Nb* = 0.03‐0.3) and depleted transition metal (e.g., Al, V contents). Most samples yield modest LREE fractionation [(La/Yb)N = 0.9‐6.1]. Nearly all samples have Chapter 4 100 positive Zr (Zr/Zr* = 0.6‐2.2) and Hf (Hf/Hf* = 0.9 ‐2.4) anomalies. LFSE‐enrichment is evident in the high Th contents relative to LREE [(Th/La)PM = 1.5‐4.4], noting one exception. Negative Eu anomalies are a feature of all felsic rocks from this formation (Eu/Eu* = 0.3‐1.0). 4.5.3.4 Chemostratigraphy of the VMS‐Host sequence Geologic sections are well constrained for the VMS‐bearing volcanic sequence of the Cerro San Lorenzo Formation and provide a framework for a chemostratigraphic model (Chap. 2; Fig 4.5). This unit‐scale evaluation permits a study of the geochemical variation within the volcanic complex and is used to investigate the potential links between the petrogenesis of the volcanic suite and generation of the massive sulphide deposits. The dacite‐rhyolite suite encompasses several geochemically distinctive units, including a syn‐mineralization rhyolite flow‐dome complex, a hanging wall dacite unit, and post‐
mineralization quartz‐phyric rhyolite dykes (Chap. 2). Basalt forms almost the entirety of the footwall and the majority of the hanging wall. SiO2 contents for the basalt units range from 45.32 to 53.48 wt.%, dacite yields 65.35 to 66.54 wt.% SiO2 and rhyolite has 68.18 to 75.16 wt.% SiO2 (Figs. 4.13A). These compositional clusters are also readily identified on the Nb/Y versus Zr/TiO2 plot (Fig. 4.13B), where basalt has low Zr/TiO2 ratios (29‐78) whereas rhyolite yields much higher ratios (534‐723) and dacite has intermediate ratios (259‐278). All samples are strongly subalkaline (Nb/Y < 0.27) with low total alkali contents dominated by Na2O (Fig. 4.13A). Mg # are substantially higher in basalt (59±8) than in dacite (46±3) or rhyolite (33±14) (Fig. 4.12C) Likewise, TiO2 contents are also distinctive for each group, ranging from 0.24‐0.43 wt.% in rhyolite, 0.58‐0.64 wt.% in dacite, and 0.70‐1.34 wt.% in basalt (Fig. 4.13D). In contrast, P2O5 contents overlap the three groups (Fig. 4.13E). Dacite also has intermediate contents of transition metals such as Co, Cr, Ni, and V. Overall, the rocks display a systematic decrease of the transition elements with more evolved compositions Chapter 4 101 (Fig. 4.13F). Within the basalt data, TiO2 and P2O5 yield a positive linear correlation with Zr and V displays a negative correlation with Zr/TiO2. Although these trends suggest fractionation within the basaltic endmembers, no fractionation trend exists between the mafic and felsic data. HFSE systematics indicates complexity in the petrogenesis of the felsic volcanic rocks. On a Hf‐Nb plot (Fig. 4.13G), basalt has a strong positive linear correlation, probably as a result of derivation from a single source, whereas dacite and rhyolite plot with higher Hf‐Nb contents and show no linear relationship. On a Zr/Y versus Y plot (Fig. 4.13H) the data define three distinct clusters, with basalt defining with the lowest Zr/Y and Y values. Syn‐mineralization rhyolite from TG1 and TG3 is characterized by relatively high Y contents and moderate Zr/Y. Post‐mineralization dacite flows and rhyolite dykes at TG1‐TG3, as well as syn‐mineralization rhyolite from B5, are characterized by moderate Y contents and the highest Zr/Y values. These groupings are mimicked on a La/YbN versus YbN plot (Fig.13I). Finally, a plot of Th versus Zr (Fig. 4.13J) illustrates that rhyolite dykes are more strongly enriched in LFSE (i.e., Th) but yield comparable HFSE (i.e., Zr) with respect to the other volcanic units. The non‐systematic clustering of felsic rocks based on variable REE, HFSE and LFSE abundances and ratios, indicate variations that do not relate to fractionation and require heterogeneity of the felsic melt source. 4.6 Discussion 4.6.1 Petrochemical Variations in Mafic Volcanic Rocks of the Lancones Basin Basalt geochemical data from the various formations within the Lancones basin indicate broad compositional similarities which are consistent with an arc volcanic setting, though each formation has distinctive geochemical characteristics. The HFSE and LFSE systematics of basalt from all formations in the Lancones basin have low Nb contents and variable Th, Zr, and Y Chapter 4 102 values. These low‐moderate HFSE, moderate REE, and variable LFSE contents are akin to those of basalt of oceanic arc affinity (Ewart and Hawkesworth, 1987; McCulloch and Gamble, 1991; Pearce and Parkinson, 1993; Pearce and Peate, 1995). Consistent with a transitional to calc‐
alkaline affinity for the Cerro San Lorenzo Formation, lower V values are indicative of a relatively oxidized magma source (i.e., higher ƒO2; Shervais, 1982). These geochemical variations are reemphasized on a Ta/Yb versus Th/Yb plot (Fig. 4.14) which suggests that basalt from the Cerro San Lorenzo Formation is related to subduction zone enrichment of an N‐MORB type mantle source producing mildly calc‐alkaline oceanic arc basalt (Pearce, 1983). Rocks from the Cerro El Ereo Formation have similar LFSE (i.e., Th) enrichment with respect to a mantle source, but are derived from a source more depleted than a MORB source, typical of tholeiitic oceanic arc basalt (Fig. 4.14). Cerro San Lorenzo Formation basalt displays higher mean (La/Yb)N but similar YbN values when compared to basalt from the Cerro El Ereo and La Bocana formations (Fig. 4.10A). HFSE ratios such as Zr/Y mimic the LREE/HREE ratios with the highest Zr/Y values occurring in basalt for the Cerro San Lorenzo Formation (Fig. 4.10B). Lower LREE/HREE and Zr/Y ratios of the La Bocana and Cerro El Ereo Formations are typical of tholeiitic rocks, whereas the higher ratios of the Cerro San Lorenzo Formation are characteristic of transitional to mildly calc‐alkaline affinity (cf. Barrett and MacLean, 1999). REE (e.g., La/Yb) and HFSE (e.g., Zr/Y) ratios, as well as conventional AFM diagrams suggest basalt from the Cerro San Lorenzo Formation is more akin to calc‐alkalic suites, whereas Cerro El Ereo and La Bocana formations are mostly island arc tholeiitic basalts. The variance in both HFSE and REE suggests heterogeneity in the source region for these rocks. The younger Cerro El Ereo and La Bocana Formations were generated from a source more depleted than that of the Cerro San Lorenzo Formation. One possible explanation is Chapter 4 103 related to the initial extraction of a melt to form the Cerro San Lorenzo Formation lavas which would have caused some depletion of incompatible elements from the mantle‐wedge source. Subsequent lavas from the same region of the mantle would have been relatively depleted. Alternatively, the geochemical variations may be explained by variations in the melt generation processes, whereby the earlier and mildly calc‐alkaline Cerro San Lorenzo Formation basalt formed through lower amounts of partial melting and slightly higher ƒO2 when compared to the younger formations (Shervais, 1982). Phase 1 basalt is characterized by MgO, Ni and Cr contents that are on average significantly higher than for basalt from the phase 2 sequence, indicating a more primitive magma source for these early‐rift lavas. ‘Primitive’ arc magmas are those with up to 8 wt% MgO in oceanic arcs, and lower values for mature arcs (Leate and Larter, 2003). Cerro San Lorenzo Formation basalt averages 7.4 ±4.4 wt% MgO, compared to 4.7 ±1.8 wt% and 3.9 wt% MgO in the younger Cerro El Ereo and La Bocana formations, respectively. The primitive lavas represented by the Cerro San Lorenzo Formation were likely erupted through relatively thin, juvenile arc crust. The more fractionated basaltic‐andesite compositions of Phase 2 formed in a maturing arc with thicker crust (Green, 1980; Gill, 1981) during the waning stages of arc rifting. The data suggest phase 1 was a period of maximum extension, rifting, and subsidence, with relatively thin arc crust and high heat flow, and therefore conducive to VMS formation. 4.6.2 Felsic Volcanic Rock Petrochemistry and Association with VMS Dacitic to rhyolitic rocks occur in the Cerro San Lorenzo and La Bocana Formations, but are absent in the Cerro El Ereo Formation. Felsic volcanic rocks from these formations are almost all subalkaline and show broadly similar patterns on normalized trace element plots. Pronounced negative Ti and Nb anomalies are ubiquitous and transition metals (e.g., Al, Sc, and V) are also strongly depleted. With minor exceptions, nearly all samples have positive Zr and Hf Chapter 4 104 anomalies. LFSE‐enrichment is evident in the high Th contents relative to LREE and HFSE. Weak to moderate negative Eu anomalies and moderate LREE/HREE ratios are characteristic of nearly all samples. There are slight variations in mean Zr/TiO2 ratios, with ratios slightly higher in the Cerro San Lorenzo Formation (674 ±314) than those of the La Bocana Formation (422 ±341). Low HFSE contents for felsic rocks reflect M‐ to I‐type magmatic affinities. In general, M‐ to I‐type felsic rocks generally reflect either a direct mantle source or are the result of partial melting of juvenile crust due to heat from a subcrustal underplate, possibly with some influence from the continental crust (Pitcher, 1983; Whalen, 1985). In a mantle source scenario, the felsic rocks were likely derived via fractionation of an arc basaltic melt, though the general lack of abundant intermediate compositions suggests this is not likely the case. Partial melting of subcrustal or crustal material which forms the arc is considered more likely due to the general bimodality of the sequence. Considering the inherited Proterozoic zircons (Chap. 2) and evolved Nd and Pb isotope signatures of felsic rocks (Chap. 5), there is strong evidence for some influence of continental crust either at the source of the melt and/or during magma ascent. It is inferred that syn‐mineralization rhyolite, which is the first erupted unit of the felsic suite and which has the least incompatible element enriched signatures, had a shorter crustal residence time than either hanging wall dacite or post‐mineral rhyolite dykes. Much work has been done on assessing the geochemistry of felsic volcanic rocks associated with VMS deposits in order to understand the petrogenesis of the felsic magmas in relation to the genesis of the deposits (e.g., Lesher et al., 1986; Barrie et al., 1993; Lentz, 1998; Hart et al., 2004; Piercey, 2007). Archean rhyolites are classified into four suites according to HFSE and REE systematics which are correlated with the relative abundance (and size) of VMS deposits (Fig. 4.16 A,B; Lesher et al., 1986). FI alkalic dacites‐rhyodacites yield the lowest HFSE and HREE Chapter 4 105 (e.g., Y, Yb) and the highest LREE/HREE (La/Yb) and Zr/Y ratios, are abundant in the rock record but are typically barren of VMS deposits (Fig. 4.16). FIII tholeiitic rhyolites have the highest Y and Yb and the lowest La/Yb and Zr/Y ratios and host the greatest number of deposits. FII rhyolites have intermediate compositions and also host VMS deposits, but are mostly barren. For some Phanerozoic and Proterozoic VMS‐associated rhyolites, a class of relatively more Y and Yb‐depleted rhyolite is defined as FIV (Hart et al., 2004). These compositions approximate partial melting of hydrous mafic crust or amphibolite at relatively high temperature and low pressure conditions, in equilibrium with hornblende (clinopyroxene at higher temperatures) and plagioclase (Beard and Lofgren, 1991; Wolf and Wyllie, 1994). As the REE are moderately to strongly incompatible in this phase assemblage, REE abundances tend to be relatively high in the melt, though strong partitioning of Eu into plagioclase is likely due to low ƒO2. FII rhyolites appear to the dominant VMS‐bearing class in the Phanerozoic and Proterozoic, not because of differences in the generation of VMS systems, but because of the evolution of plate tectonics and variation in petrogenesis of felsic magmas in these terrains (Lesher et al., 1986; Lentz, 1998). Piercey (2007) illustrates that the petrochemical classification of rhyolites in post‐Archean terrains is dependent on whether the magmas were generated in juvenile or evolved environments. In post‐Archean evolved terrains, most VMS‐associated rhyolites are calc‐alkalic FII‐FIII, though some are associated with FI (Piercey, 2007). In post‐Archean juvenile terrains, VMS deposits are associated with tholeiitic, low La/Yb, Zr/Y ratios and strongly HFSE‐depleted rhyolites that mostly overlap FII and FIV. Despite time‐dependant geochemical variations, the consensus amongst researchers is that VMS‐associated magmas are high temperature (>900o) shallow crustal (<10 km) melts of either hydrated mafic crust (in Archean and post‐Archean juvenile terrain) or continental crust (post‐
Chapter 4 106 Archean evolved settings), whereas non‐VMS rhyolites are generated at deeper crustal levels and possibly lower temperatures (Lesher et al., 1986; Barrie et al., 1993; Hart et al., 2004; Piercey et al., 2001). High temperature melting of crust at shallow depths imply a high geothermal gradient, a thinned crust and extensional tectonics which favour VMS hydrothermal systems due to high heat flow and permeability (Lesher et al., 1986; Hart et al., 2004). VMS‐associated felsic volcanic rocks at Tambogrande (Cerro San Lorenzo Formation) yield relatively low to moderate (La/Yb)N and Zr/Y, low Yb, Y and Zr, and moderate‐strong negative Eu anomalies. These values overlap with FII to FIII rhyolites (Fig. 4.16 A,B; Lesher et al., 1986; Hart et al., 2004) and other Archean Superior Province VMS‐associated rhyolites (e.g., group II of Barrie et al., 1993; Fig. 4.16 C‐D). VMS‐barren felsic volcanic rocks from the La Bocana Formation have similar but slightly more depleted HFSE and REE and mostly overlap the FIV group. Compared to some Phanerozoic VMS districts (Lentz, 1998), the Tambogrande data resemble the Mt. Windsor and Kuroko districts, examples of a back‐arc or intra‐arc rift settings with continental basement (Fig. 4.16E‐F). Considering the inherited Proterozoic zircons (Chap. 2) and evolved Nd and Pb isotope signatures of felsic rocks at Tambogrande (Chap. 5), there is strong evidence for some influence of continental crust either at the source of the melt and/or during magma ascent. It is inferred that syn‐mineralization rhyolite, which is the first erupted unit of the felsic suite and which has the least incompatible element enriched signatures, had a shorter crustal residence time than either hanging wall dacite or post‐mineralization rhyolite dykes. These high temperature and aphyric (i.e., at liquidus) felsic magmas ascended rapidly from the area of melt generation in the middle to lower crust and were erupted in the vicinity of Chapter 4 107 hydrothermal venting and VMS formation on the seafloor. Some models propose that a rapid ascent of high‐temperate felsic magma to surficial environment enhances VMS formation by driving the hydrothermal system (e.g., Barrie et al., 1999. However, in the VMS genesis model proposed in Chapter 5, felsic volcanic rocks have a passive genetic role in the formation of VMS deposits at Tambogrande. Rather, high temperature felsic volcanism is symptomatic of the anomalous tectonic and geothermal setting. This is supported by the Pb isotope data which yield distinctively different isotopic signatures for massive sulphide mineralization and felsic volcanic rock (Chap. 5). 4.6.3 Implications for the Tectonic Setting The petrochemical volcanic rock assemblages for the Lancones basin suggest oceanic island arc geochemical affinities. For the generation of the mafic volcanic rocks, partial melting of a depleted mantle or MORB‐type source (i.e., the mantle wedge) with the influence of a slab‐
derived ‘fluid’ component is inferred and is the principal process accepted in arc volcanic settings for the genesis of basalts of arc tholeiitic to calc‐alkalic affinity (McCulloch and Gamble, 1991; Hawkesworth et al., 1993; Pearce and Parkinson, 1993; Pearce and Peate, 1995). According to HFSE systematics, felsic volcanic rocks from the Lancones Basin are dominantly M‐type (direct mantle source) transitional to I‐type (mafic, indirect mantle‐derived source, e.g., subcrustal underplate), typical of oceanic island arc terrains (Pitcher, 1983; Whalen, 1985). Geochemical data, such as variable HFSE (Zr, Y, Nb), LFSE (Th) and REE implicate continental crust in the generation of the source of the felsic volcanic rocks. Therefore, the volcanic arc was developed as a ‘continental margin arc‐rift’ but resembles many modern oceanic island arc rifts or possible nascent back‐arcs. In order to generate oceanic arc petrochemical affinities for volcanic rocks in a continental margin setting, crustal thickness of the overriding plate must have been sufficiently thin to Chapter 4 108 allow for limited modification of ascending mantle‐derived magmas. This was likely the case along the western Peruvian margin during the Cretaceous when the Andean orogeny was in the incipient stages of development and the leading edge of the continental crust was much thinner than at present (Benavides‐Cáceres, 1999). Moreover, the newly developed suture bounding the Amotape allocthon during the Early Cretaceous likely provided a zone of crustal weakness along which rifting occurred. A composite margin with inherent structural weaknesses would have more easily enabled crustal attenuation and rifting than a coherent zone of uniform crustal composition. Given that back‐arc basin basalts can show a geochemically continuity with island arc tholeiites (Taylor and Martinez, 2003), the recognition of the back‐arc in ancient arc settings is a somewhat difficult task. Unless MORB‐type rocks are present, which imply well established spreading away from the arc, back arc processes are difficult to verify. Geochemical and geologic data from the Lancones basin suggest that an extensional arc regime (i.e., intra‐arc rift or nascent back‐arc) was responsible for the development of a marginal basin in northwestern Perú during the Cretaceous. No ‘non‐arc’ volcanic rocks (e.g., MORB) were identified which might indicate intra‐continental rifting or back arc spreading within the Lancones basin. In the Western Peruvian Trough, however, Atherton et al. (1985) suggest some basalt may have MORB geochemical characteristics related to incipient back‐arc spreading. However, Cretaceous intra‐continental alkaline volcanic rocks are preserved in the Eastern Cordillera from Ecuador to Argentina and represent the associated back‐arc (Soler and Bonhomme, 1990; Viramonte et al., 1999; Barragán et al., 2005). Therefore, a rifted arc (± incipient back‐arc spreading) in a peri‐continental setting is an acceptable analogue to the Lancones basin. Chapter 4 109 Extensional basins such as Bransfield Strait (Barker and Austin, 1998) and Manus Basin (Sinton et al., 2003) are analogous modern settings. 4.7 Summary In summary, there are several petrochemical features of the early rift stage volcanism (i.e., Cerro San Lorenzo Formation) that are indicative of favourable geologic conditions for VMS formation. The presence of petrochemically distinct syn‐ and post‐mineralization felsic volcanic units suggests a complex igneous system with multiple magma batches. This favours a thermally and seismically dynamic setting required to sustained hydrothermal circulation and VMS formation. High MgO mafic volcanic rocks of the Cerro San Lorenzo Formation are the most primitive of all volcanic rocks in the Lancones basin. Moreover, associated with the VMS deposits are M‐
type, FII‐FIII rhyolite, considered to be a high temperature partial melt of juvenile crustal derivation. These bimodal volcanic rocks retained much of the heat of fusion through to eruption and therefore indicate a high temperature gradient in the crust, conducive to VMS formation. Phase I volcanism was associated was a rifting event along the continental margin. However, subsequent phase II volcanism is derived from a more depleted mantle (recycling of mantle wedge without advection of ‘new’ mantle) and yields a more fractionated mafic endmember. Coupled with the greater abundance of felsic volcanic rocks from a more depleted source, a waning extensional tectonic regime is inferred. The more robust extension during phase 1 favours a high geothermal gradient and enhanced hydrothermal circulation along syn‐volcanic faults, components integral to VMS formation. Phase 2 appears to be a tectonomagmatic stage that is transitional to granitic plutonism of the Coastal batholith. Chapter 4 110 Figure 4.1. Morphostructural units of the Peruvian Andes (modified after Benavides‐Cáceres, 1999). Cretaceous marginal basins ‐ Lancones (LB), Huarmey (HB) and Cañete (CB) basins ‐ are superimposed. Also shown are the locations of VMS deposits and prospects (circles) (data from Steinmüller et al., 2000). Chapter 4 111 Figure 4.2. A. Location map for the Tambogrande project; B. regional map showing major tectonostratigraphic units of coastal northwestern Perú. The locations of VMS deposits (TG1, TG3, and B5) in the Tambogrande area are also shown and field area of this study outlined (see Fig. 4.3 for a detailed map). Modified after Jaillard et al. (1999), Tegart et al. (2000). Chapter 4 112 Figure 4.3 – Location map and simplified cross sections along the Peruvian continental margin based on gravity modeling and seismic data from Couch et al. (1981) and Jones (1981). Chapter 4 113 Figure 4.4 (following page). Regional geologic map for the Tambogrande area of the Lancones basin reviewed in this study. The location of VMS deposits TG1, TG3, and B5, as well as geochemical sampling locations are shown. Map projection is WGS 84, Zone 17S. Map is from this study. Chapter 4 114 Chapter 4 115 Figure 4.5. Schematic stratigraphic column of the volcanic arc sequence of the Lancones basin. Inset section is a more detailed schematic section of the VMS‐bearing sequence at Tambogrande. Chapter 4 116 Figure 4.6. A. Silica vs. total alkalies classification scheme of Le Bas et al. (1986). B. Nb/Y versus Zr/TiO2 plot of Winchester and Floyd (1977). C. AFM plot (Irvine and Baragar, 1971). Chapter 4 117 Figure 4.7. Bivariate plots of basalt from the Cerro San Lorenzo, Cerro El Ereo and La Bocana formations. A. SiO2 versus MgO, B. Ni versus MgO, C. Cr versus MgO, and D. V versus MgO. The vertical dashed line at 5.5 w% MgO emphasizes the division in the data. Chapter 4 118 Figure 4.8. Basalt discrimination diagrams. A. Th‐Zr‐Nb plot (Wood, 1980). B. Zr‐Nb‐Y plot (Meschede, 1986); all samples illustrate relatively low Nb but variable Th, Zr, and Y values and are defined as arc basalt. C. Zr‐Ti‐Y plot (Pearce and Cann, 1973). D. Ti‐V plot (Shervais, 1982). Chapter 4 119 Figure 4.9. Primitive mantle‐ normalized extended trace element diagrams for mafic intermediate rocks from the various formations. A. Cerro San Lorenzo Formation basalt. B. Cerro San Lorenzo Formation basaltic‐andesite. C. Cerro El Ereo Formation basalt. D. Cerro El Ereo Formation basaltic‐andesite. D. La Bocana Formation basalt. E. La Bocana Formation basaltic‐andesite. Element order and normalizing values follow Sun and McDonough (1989). Chapter 4 120 Figure 4.10. Chondrite‐normalized (using values from Sun and McDonough, 1989) HFSE values for basalt from the Cerro San Lorenzo, Cerro El Ereo, and La Bocana formations. A. Yb versus La/Yb. B. Y versus Zr/Y. Chapter 4 121 Figure 4.11. Felsic volcanic discrimination diagrams. A. Ga/Al versus Zr after Whalen et al. (1987). B. Y versus Nb (Pearce et al., 1984). All samples plot within the I‐ to M‐type field. Chapter 4 122 Figure 4.12. Primitive mantle‐normalized extended trace element diagrams show broadly similar patterns for felsic volcanic rocks of the Lancones basin. A. Cerro San Lorenzo Formation. B. La Bocana Formation. Inset diagrams are rare earth elements only and are normalized to chondrite values. Element order and normalizing values follow Sun and McDonough (1989). Figure 4.13 (following page). Multi‐element bivariate plots for basalt, dacite, and rhyolite in the vicinity of VMS deposits at Tambogrande. A. Zr/TiO2 vs. Nb/Y (Winchester and Floyd, 1977). B. Na2O+K2O vs. SiO2 (Le Bas et al., 1986). C. Fe2O3+MgO vs. SiO2. D. TiO2 vs. Zr. E. P2O5 vs. Zr. F. V vs. Zr/TiO2. G. Hf vs. Nb. H. Zr/Y vs. Y. I. La/Yb vs Yb (chondrite‐normalized values using Sun and McDonough (1989)). J. Th vs. Zr. Chapter 4 123 Chapter 4 124 Figure 4.14. Ta/Yb versus Th/Yb plot after Pearce (1983). Only Cerro San Lorenzo Formation and El Ereo Formation basalt shown due to data limits for La Bocana Formation. D = depleted mantle. E = enriched mantle. Chapter 4 125 Figure 4.15. Schematic two‐stage tectonic model for arc magmatism in the Lancones basin. The model proposes a shift from phase 1, extensional tectonics with a steeply dipping subduction zone to phase 2, waning extension and shallower subduction. Depletion of the mantle‐wedge explains the relatively HFSE‐ and REE‐depleted mafic volcanic rocks in the phase 2. The thickened crust due to waning extension forces more fractionation of mafic intrusions. The partial melting of the arc crust causes phase 2 felsic volcanic rocks to yield lower HFSE contents. Chapter 4 126 Figure 4. 16. Felsic volcanic rock discrimination diagrams (after Lesher et al., 1986; Barrie et al., 1993; Lentz, 1998; Hart et al., 2004). A. (La/Yb)N versus YbN (Chondrite‐normalized using values from Sun and McDonough (1989)). B. Zr/Y versus Y. Chapter 4 127 Table 4.1 (Following page). Average lithogeochemical values and 2 σ errors for volcanic rocks of the Lancones basin, based on volcanic rock type and formation. Complete data are listed in Appendix C. Chapter 4 128 SiO2 (wt.%)
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2 O
P2O5
LOI
TOTAL
Cr (ppm)
Ni
Co
Sc
V
Li
Cu
Pb
Zn
Rb
Cs
Ba
Sr
Ga
Ta
Nb
Hf
Zr
Y
Th
U
Chapter 4 CSLF
basalt
(n=42)
49.62
0.90
16.20
9.54
0.16
7.40
8.37
3.02
0.66
0.15
3.59
99.67
227
68
36
33
308
8
89
7
110
16
1
498
225
17.0
0.1
2.0
1.6
53
17
2.1
0.7
2σ
4.67
0.48
2.69
2.61
0.10
4.38
6.96
2.36
1.25
0.15
4.83
1.38
CEEF
basalt
(n=7)
48.91
0.74
18.67
10.54
0.23
4.74
10.71
2.60
0.31
0.06
2.23
99.76
2σ
3.14
0.34
1.81
1.96
0.29
1.83
6.08
2.34
0.50
0.04
2.54
1.21
566
121
19
11
154
9
146
7
116
31
1
1251
178
4.9
0.4
2.3
1.1
34
9
2.9
0.9
79
19
30
40
360
6
109
5
83
6
1
156
193
17.4
0.0
0.8
1.3
30
15
0.5
0.5
94
21
10
14
54
0
156
14
50
10
0
157
78
1.0
0.0
0.3
0.4
24
8
0.7
0.0
LBF
basalt
(n=9)
51.43
0.83
16.80
10.59
0.18
3.92
9.76
3.00
0.39
0.12
2.99
100.05
2σ
4.88
0.44
5.10
5.44
0.06
1.46
7.43
2.20
0.59
0.11
5.11
0.63
129
24
33
416
34
13
434
n/a
159
6
100
8
1
235
224
18.8
n/a
2.0
1.5
39
21
2.3
0.8
259
127
4
60
14
1
350
162
4.8
2.0
1.5
49
12
3.8
0.9
CSLF
basalticandesite
(n=10)
58.28
0.82
16.30
6.97
0.13
3.08
4.82
4.93
1.01
0.09
2.71
99.40
61
15
19
16
151
3
71
5
87
2
1
758
255
5
0.4
4.7
4.0
130.6
28
7
1.8
2σ
3.29
3.29
0.36
1.90
2.86
0.07
1.75
3.07
1.74
0.77
0.22
3.55
CEEF
basalticandesite
(n=1)
59.12
0.97
15.32
9.93
0.18
2.08
6.70
4.54
0.18
0.15
1.11
100.30
LBF
basalticandesite
(n=10)
61.92
0.63
15.74
6.33
0.13
1.53
5.47
4.63
0.72
0.17
2.71
100.06
16
50
17
19
7
127
6
140
7
55
15
2
1177
200
3.3
0.2
3.4
1.8
58
15
5.3
90
5
18
n/a
145
n/a
20
5
100
2
0
121
101
22.0
<0.5
1.0
1.0
58
29
1.0
0.5
86
14
10
n/a
126
n/a
61
5
63
22
1
520
352
14.6
0.5
3.3
3.1
109
23
4.8
1.4
114
0
41
33
2
730
240
6.3
0.0
4.4
2.1
78
24
4.2
1.3
2σ
5.63
0.27
3.73
4.82
0.14
1.42
2.79
2.11
1.28
0.08
1.97
0.55
55
14
7
121
CSLF
postmineralization
andesite
(n=7)
67.26
0.63
14.47
4.11
0.07
1.62
2.96
4.85
1.38
0.17
2.37
99.94
81
11
12
15
51
2
29
3
52
24
1
788
142
16.3
0.2
4.1
4.9
160
34
5.2
2.2
2σ
5.67
0.11
1.79
1.93
0.05
1.38
3.69
2.29
2.17
0.08
2.42
0.60
CSLF
synmineralization
rhyolite
(n=22)
71.68
0.30
13.18
2.99
0.05
1.13
1.59
5.35
1.03
0.07
2.00
99.44
2σ
6.44
0.15
2.44
2.31
0.06
1.73
1.44
2.84
1.87
0.06
2.31
1.65
66
14
12
2
45
5
27
8
35
40
2
1131
190
6.2
0.1
1.2
1.5
32
13
3.4
1.2
125
6
4
10
15
3
56
2
76
21
1
642
99
16.7
0.4
5.4
5.6
195
39
7.3
3.0
152
9
7
12
32
4
181
8
116
41
1
1739
111
4.1
0.3
5.9
2.2
79
23
8.2
2.8
129 Table 4.1. Continued SiO2 (wt.%)
TiO2
Al2O3
Fe2O3
MnO
MgO
CaO
Na2O
K2O
P2O5
LOI
TOTAL
Cr (ppm)
Ni
Co
Sc
V
Li
Cu
Pb
Zn
Rb
Cs
Ba
Sr
Ga
Ta
Nb
Hf
Zr
Y
Th
U
Chapter 4 CSLF postmineralization
rhyolite
dyke (n=3)
71.79
0.31
13.17
3.13
0.08
0.96
1.89
4.50
1.06
0.06
2.67
99.63
2σ
5.13
0.17
0.94
1.26
0.03
0.33
1.43
1.04
0.86
0.03
1.62
0.88
36
0
5
8
19
1
10
0
83
25
16
4
4
3
19
4
32
6
12
19
764
127
n/a
0.3
3.7
5.4
177
31
8.5
n/a
906
20
0.1
1.9
0.9
37
7
0.6
LBF
daciterhyolite
(n=17)
72.66
0.31
13.01
2.77
0.07
0.86
2.13
4.91
0.84
0.07
2.04
99.74
2σ
4.36
0.10
1.37
1.14
0.06
0.80
1.08
1.22
0.74
0.04
1.46
0.44
121
12
4
n/a
23
n/a
46
8
63
17
1
518
120
14.0
0.5
2.4
3.7
117
33
4.5
2.1
65
7
2
17
57
4
42
17
0
429
62
3.0
0.0
1.9
1.2
42
10
3.2
1.5
130 Table 4.1. Continued La (ppm)
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
CSLF
basalt
(n=42)
6.8
14.8
2.1
9.7
2.7
0.9
2.9
0.5
3.0
0.7
2.0
0.3
1.9
0.3
2σ
CEEF
basalt
(n=7)
7.0
13.5
1.7
7.1
1.6
0.5
1.5
0.2
1.5
0.3
1.0
0.2
0.9
0.1
2.6
5.8
0.9
4.8
1.6
0.6
2.0
0.4
2.5
0.6
1.7
0.2
1.7
0.2
2σ
LBF
basalt
(n=9)
3.1
4.4
0.7
3.5
1.0
0.3
1.2
0.2
1.4
0.3
1.0
0.1
0.8
0.2
4.2
8.8
1.3
6.6
2.2
0.8
2.6
0.5
3.3
0.8
2.2
0.4
2.3
0.4
2σ
CSLF
basalticandesite
(n=10)
7.1
14.4
1.9
8.6
2.1
0.5
1.8
0.3
1.9
0.5
1.3
0.2
1.2
0.2
15.6
34.2
4.7
20.1
4.9
1.4
4.9
0.8
4.8
1.1
3.2
0.5
3.2
0.5
2σ
CEEF
basalticandesite
(n=1)
LBF
basalticandesite
(n=10)
2σ
CSLF
postmineral
-ization
andesite
(n=7)
1.7
10.8
22.1
2.8
11.7
2.5
0.8
2.3
0.4
2.4
0.5
1.6
0.2
1.4
6.0
14.5
2.2
11.0
3.5
1.1
3.7
0.7
4.6
1.1
2.9
0.5
3.2
0.5
14.3
28.8
3.8
15.4
3.8
1.0
4.0
0.7
3.8
0.9
2.4
0.4
2.4
0.4
10.0
22.8
3.1
12.2
3.3
0.5
3.5
0.6
3.8
0.9
2.5
0.4
2.6
0.4
15.1
32.8
4.5
19.5
5.1
1.4
5.5
0.9
5.8
1.4
4.2
0.6
4.0
0.6
2σ
CSLF
synmineral
-ization
rhyolite
(n=22)
2σ
12.1
22.6
3.0
9.9
1.8
0.6
1.6
0.4
1.7
0.5
1.4
0.2
1.2
0.2
17.0
37.4
5.0
21.5
5.5
1.3
5.9
1.0
6.5
1.5
4.6
0.7
4.5
0.7
10.4
19.5
2.3
8.9
2.4
0.6
2.7
0.5
3.7
0.8
2.5
0.4
2.3
0.4
Chapter 4 131 Table 4.1. Continued La
Ce
Pr
Nd
Sm
Eu
Gd
Tb
Dy
Ho
Er
Tm
Yb
Lu
CSLF postmineralization
rhyolite dyke
(n=3)
18.3
37.5
4.8
19.7
4.8
1.2
4.6
0.8
5.2
1.2
3.9
0.6
3.8
0.6
2σ
LBF
daciterhyolite
(n=17)
2σ
2.7
6.0
0.8
4.1
1.1
0.5
1.1
0.2
1.2
0.3
0.8
0.1
0.7
0.1
11.2
23.8
3.3
14.5
4.0
0.9
4.4
0.8
5.2
1.2
3.5
0.5
3.6
0.6
4.8
10.8
1.4
5.4
1.2
0.3
1.3
0.2
1.6
0.4
1.1
0.2
1.2
0.2
Abbreviations: CSLF = Cerro San Lorenzo Formation; CEEF = Cerro El Ereo Formation; LBF = La Bocana Formation. Chapter 4 132 Table 4.2. Summary of average selected trace element ratios for volcanic rocks of the Lancones basin based on volcanic rock type and formation. CSLF
basalt
(n=42)
2σ
CEEF
basalt
(n=7)
2σ
LBF
basalt
(n=9)
2σ
CSLF
basalticandesite
(n=10)
2σ
Nb/Y
Nb/Yb
Zr/Nb
La/Nb
Th/Nb
0.1
1.1
31.0
3.8
1.1
0.1
1.2
19.4
3.4
1.3
0.0
0.4
57.7
4.1
0.4
0.0
0.0
10.8
3.8
0.1
0.1
0.6
45.7
5.8
1.3
0.1
0.6
23.7
3.3
0.7
1.5
29.2
3.5
1.5
0.1
1.0
10.8
2.4
(Th/La)PM
Zr/TiO2
Zr/Y
(La/Yb)N
Eu/Eu*
Nb/Nb*
Zr/Zr*
Hf/Hf*
Ti/Ti*
2.3
59
3.1
2.5
0.98
0.19
0.86
0.96
0.79
1.8
31
1.7
2.7
0.17
0.18
0.34
0.55
0.37
1.0
40
2.0
1.0
1.07
0.26
0.78
1.07
1.03
2.0
17
0.9
0.7
0.17
0.20
0.47
0.89
0.55
2.2
47
1.8
1.1
1.00
0.12
0.71
0.89
0.88
1.8
40
1.1
1.4
0.23
0.01
0.24
0.28
0.48
2.5
164
4.8
42.4
0.83
0.17
1.25
1.40
0.39
1.4
5
65.7
2.0
2.41
0.29
0.17
0.60
0.84
CSLF
postmineralization
andesit
e (n=7)
CSLF
synmineral
-ization
rhyolite
(n=22)
CEEF
basalticandesite
(n=1)
LBF
basalticandesite
(n=10)
0.0
0.3
58.0
6.0
0.1
1.3
47.8
6.0
1.9
0.2
1.7
59.9
5.5
2.1
0.1
1.0
39.6
3.6
1.2
0.1
1.0
15.6
2.4
0.6
0.1
1.2
43.9
3.8
1.6
0.3
2.5
33.9
4.2
1.9
0.0
60
2.0
1.3
0.93
0.14
0.74
0.46
0.63
2.7
184
5.3
4.6
0.86
0.11
1.35
1.39
0.39
2.1
124
4.6
3.9
0.28
0.10
0.89
0.81
0.17
3.1
255
4.9
2.6
0.78
0.16
1.42
1.60
0.29
2.4
35
2.2
2.0
0.16
0.07
0.66
0.87
0.11
2.7
674
5.4
2.8
0.68
0.17
1.60
1.66
0.13
1.5
314
3.6
2.7
0.23
0.12
0.65
0.65
0.09
2σ
2σ
Chapter 4 133 2σ
Table 4.. Continued 2σ
LBF
daciterhyolite
(n=17)
2σ
0.1
1.0
49.6
5.2
2.4
0.3
2.7
17.4
2.1
1.2
0.1
0.7
69.3
6.6
2.2
0.2
1.7
31.7
3.6
1.3
3.8
594
5.7
3.3
0.7
0.1
1.7
1.9
0.2
0.5
246
0.2
0.2
0.1
0.0
0.1
0.2
0.1
2.8
422
3.7
2.3
0.7
0.1
1.3
1.5
0.2
1.0
171
1.2
1.3
0.2
0.1
0.4
0.4
0.1
CSLF postore rhyolite
dyke (n=3)
Nb/Y
Nb/Yb
Zr/Nb
La/Nb
Th/Nb
(Th/La)PM
Zr/TiO2
Zr/Y
(La/Yb)N
Eu/Eu*
Nb/Nb*
Zr/Zr*
Hf/Hf*
Ti/Ti*
Notes: CSLF, CEEF, LBF as per table 1. PM = primitive mantle normalized N = chondrite normalized Eu/Eu* = 0.5EuPM /(GdPM + SmPM) Nb/Nb* = 0.5*NbPM/(ThPM + LaPM) Zr/Zr* = 0.5*ZrPM/(GdPM + SmPM) Hf/Hf* = 0.5*HfPM/(GdPM + SmPM) Ti/Ti* = 0.5*TiPM/(GdPM + SmPM) Chapter 4 134 4.8 References Aguirre L, Offler, R. 1985. Burial metamorphism in the Western Peruvian trough: its relation to Andean magmatism and tectonics. Edited by W.S. Pitcher, M.P. Atherton, E.J. Cobbing, and R.D. Beckinsdale. In Magmatism at a plate edge: the Peruvian Andes, Blackie, Glasgow, pp. 59–71. Arculus, R.J., Lapierre, H., and Jaillard, E. 1999. Geochemical window into subduction and accretion processes; Raspas metamorphic complex, Ecuador. Geology, 27: 547‐550. Aspden, J.A., Bonilla, W. and Duque, P., 1995. The El Oro metamorphic complex, Ecuador: geology and economic mineral deposits. British Geological Survey, Overseas Geology and Mineral Resources Series 67, 63 p. Atherton, M.P., 1990. The Coastal Batholith of Peru: the product of rapid recycling of new crust formed within rifted continental margin. Geological Journal, 25: 335–349. Atherton, M.P., Pitcher, W.S., Warden, V. 1983. The Mesozoic marginal basin of central Perú. Nature, 350: 303‐306. Bailes, A.H. and Galley, A.G. 1999. Evolution of the Paleoproterozoic Snow Lake arc assemblage and geodynamic setting for associated volcanic‐hosted massive sulphide deposits, Flin Flon belt, Manitoba, Canada. Canadian Journal of Earth Sciences, 36: 1789‐1805. Barker, D. H. N., and Austin, Jr. J. A. 1998. Rift propagation, detachment faulting, and associated magmatism in Bransfield Strait, Antarctic Peninsula, Journal of Geophysical Research, 103(B10): 24,017‐24,043. Barragán, R., Babya, P., and Duncan, R. 2005. Cretaceous alkaline intra‐plate magmatism in the Ecuadorian Oriente Basin: Geochemical, geochronological and tectonic evidence. Earth and Planetary Science Letters, 236: 670‐690. Barrett, T.J. and MacLean, W.H., 1999. Volcanic sequences, lithogeochemistry, and hydrothermal alteration in some bimodal volcanic‐associated massive sulfide systems. In Volcanic‐associated massive sulphide deposits: processes and examples in modern and ancient settings. Edited by C.T. Barrie and M.D. Hannington. Reviews in Economic Geology, Volume 8, p. 101‐131. Barrie, C.T., Ludden, J.N., Green, T.H. 1993. Geochemistry of volcanic rocks associated with Cu‐
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218 Mitouard, P., Kissel, C., and Laj, C. 1990. Post‐Oligocene rotations in southern Ecuador and northern Perú and the formation of the Huancabamba deflection in the Andean Cordillera. Earth and Planetary Science Letters, 98: 329‐339. Chapter 4 Morris, R.C. and Aleman, A.R. 1975. Sedimentation and tectonics of the Middle Cretaceous Copa Sombrero Formation in Northwest Perú. Boletin de la Sociedad Geologica del Perú, 48: 49‐64. Mourier, T., Laj, C., Mégard, F., Roperch, P., Mitouard, P., and Farfau Medrano, A. 1988. An accreted continental terrane in northwestern Perú. Earth and Planetary Science Letters, 88: 182‐192. Noble, S.R., Aspden, J.A., and Jemielita, R.A. 1997. Northern Andean crustal evolution; new U‐
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437. Sun, S. S. and McDonough, W. F. 1989. Chemical and isotopic systematics of oceanic basalts; implications for mantle composition and processes. In Magmatism in the Ocean Basins. Edited by A.D. Saunders and M.J. Norry. Geological Society Special Publication, v.42, pp. 313–345. Swinden, H.S., Jenner, G.A., and Szybinski, Z.A. 1997. Magmatic and tectonic evolution of the Cambrian‐Ordovician Laurentian margin of Iapetus; geochemical and isotopic constraints from the Notre Dame subzone, Newfoundland. In the nature of magmatism in the Appalachian orogen. Edited by K. Sinha, J.B. Whalen, and J.P. Hogan. Geological Society of America, Memoir 191, pp. 337‐365. Swinden, H.S. 1991. Paleotectonic settings of volcanogenic massive sulphide deposits in the Dunnage Zone, Newfoundland Appalachians: Canadian Institute of Mining and Metallurgy Bulletin, 83: 59–69. Chapter 4 Syme, E.C. and Bailes, A.H. 1993. Stratigraphic and tectonic setting of early Proterozoic volcanogenic massive sulfide deposits, Flin Flon, Manitoba, Economic Geology and the Bulletin of the Society of Economic Geologists, 88: 566‐589. Taylor, B., and Martinez, F. 2003. Back‐arc basin basalt systematics. Earth and Planetary Science Letters, 210: 481‐497. Tegart, P., Allen, G., and Carstensen, A., 2000, Regional setting, stratigraphy, alteration and mineralization of the Tambo Grande VMS district, Piura Department, Northern Perú. In VMS deposits of Latin America. Edited by R. Sherlock, and M.A.V. Logan. Geological Association of Canada, Mineral Deposits Division, Special Paper No. 2, pp. 375‐405. Vergara, M., Beatriz, L., Nyström, J.O., and Cancino, A. 1995. Jurassic and Early Cretaceous island arc volcanism, extension, and subsidence in the Coast Range of central Chile. Geological Society of America Bulletin, 107: 1427‐1440. Viramonte, J. G., Kay, S. M., Becchio, R., Escayola, M., and Novitski, I. 1999. Cretaceous rift related magmatism in central‐western South America. Journal of South American Earth Sciences, 12: 109‐121. Winchester, J.A. and Floyd, P.A. 1977. Geochemical discrimination of different magma series and their differentiation products using immobile elements, Chemical Geology, 20: 325‐343. Whalen, J.B. 1985. Geochemistry of an island‐arc plutonic suite: the Uasilau‐Yau Yau intrusive complex, New Britain, PNG. Journal of Petrology, 26: 603–632. Whalen, J.B., Currie, K.L., Chappell, B.W. 1987. A‐type granites: geochemical characteristics, discrimination and petrogenesis. Contributions to Mineralogy and Petrology, 95: 407–419. Winter, L.S., Tosdal, R.M., Franklin, J.M., Tegart, P. 2004. A Reconstructed Cretaceous Depositional Setting for Giant VMS Deposits at Tambogrande, Northwestern Peru. In Andean Metallogeny: New Discoveries, Concepts, and Updates. Edited By R. Sillitoe, Society of Economic Geologists Special Publication No.11, p. 319‐340. Wise, J.M. 2000. Problems with the western Peruvian trough of central Peru: an accreted oceanic arc or back‐arc rift setting? 2000 Congreso, Sociedad Geological del Peru, Lima. Wolf, M.B. and Wyllie, P.J. 1994. Dehydration‐melting of amphibolite at 10 kbar: the effect of temperature and time. Contributions to Mineralogy and Petrology, 115: 369–383. Wood, D.A. 1980. The application of the Th‐Hf‐Ta diagram to problems of tectonomagmatic classification and to establishing the nature of crustal contamination of basaltic lavas of the British Tertiary volcanic province. Earth and Planetary Science Letters, 50: 11‐30.
Chapter 4 Chapter 5. Pb‐Sr‐Nd Isotope Systematics of Cretaceous Arc Volcanic Rocks in the Lancones Basin near Tambogrande, Perú – Implications for VMS Deposit Formation 6 5.1 Overview Radiogenic (Pb‐Nd‐Sr) isotope data for volcanic rocks and sulphide minerals at the Cretaceous Cu‐Zn‐Au‐Ag volcanogenic massive sulphide (VMS) deposits at Tambogrande, Perú permit a better understanding of the tectonomagmatic setting in which the deposits formed and help constrain metal sources and models of massive sulphide formation. Based on the trace element and isotope compositions, bimodal mafic to felsic volcanic rocks at Tambogrande have a common parentage but are not related solely by fractional crystallization. Rather, magma genesis occurred at multiple stages at various crustal levels and likely involved recycling of juvenile arc crust as well as contamination with older continental crust. VMS‐associated mafic volcanic rocks yield Nd (εNdί = 4.73 – 5.93), Pb (206Pb/204Pb = 18.94 to 19.02, 207Pb/204Pb = 15.66 to 15.68), and Sr isotopic ratios (87Sr/86Srί = 0.7036‐0.7042) most akin to ocean arc settings. These isotopic values reflect binary mixing of a depleted mantle source (i.e., mantle‐wedge) and a subducted slab component consistent with the inferred tectonic setting along the western South American continental margin during the mid Cretaceous. The slab component was inherited either through subduction of continental crust‐derived sedimentary rocks and/or possibly by direct assimilation during ascent of magmas through the upper crust. Felsic volcanic rocks (εNdί = 1.85 to 5.59) are up to 4 εNd units lower, on average, than the mafic volcanic rocks, yield significantly higher 206Pb/204Pb ratios (19.64 to 20.35) and define a relatively flat Pb isotope array extending from the basalt field to a more radiogenic crustal 6
A version of this chapter will be submitted for publication. Winter, L.S. and Tosdal, R., Pb-Sr-Nd Isotope
Systematics of Cretaceous Arc Volcanic Rocks in the Lancones Basin near Tambogrande, Perú – Implications for
VMS deposit formation.
Chapter 5 Page 142 source (i.e., Olmos Complex). Sr isotope data for felsic volcanic rocks (87Sr/86Srί = 0.7045‐
0.7050) yield ratios similar but slightly greater than associated mafic rocks. The felsic volcanic rocks originated from a melt of direct basalt parentage or a partial melt of juvenile mafic crust and subsequent assimilation of radiogenic continental crust at upper crustal levels. Proterozoic to Paleozoic basement rocks and derived sedimentary successions, exposed adjacent to the Cretaceous volcanic arc in northwestern Perú, are the probable source of the crustal contamination. Pb isotopic data for each of the three VMS deposits at Tambogrande define a narrow range of values (206Pb/204Pb =18.87 to 18.93, 207Pb/204Pb = 15.67 to 15.70, 208Pb/204Pb = 38.71 to 38.85) and define steep arrays on Pb isotope plots. The data indicate a uniform source for the lead, and by inference the other metals, in these deposits, perhaps extracted during a common hydrothermal system. Hydrothermal leaching of a basalt‐dominated substrate is considered the dominant process of lead addition to the mineralizing fluids, with no apparent contribution from the more radiogenic felsic volcanic rocks or the continental crust from which the felsic rocks derive their isotopic signature. Despite derivation or modification at a shallow crustal level, felsic source melt regions and continental crust were likely beyond the deepest reaches of the hydrothermal system. Therefore, although the felsic volcanic rocks yield an important spatial and temporal association with massive sulphide deposits at Tambogrande, there is no evidence for a causal link (i.e., no leaching or magmatic fluid contribution). The presence of felsic volcanic rocks in the host rock sequence of the VMS deposits is most likely due to melting of hydrated basaltic crust and is a function of the thermal anomaly which drove the hydrothermal convection. Chapter 5 Page 143 5.2 Introduction The Tambogrande district of northwestern Perú is hosted in a Cretaceous marine volcanic and sedimentary succession in the Lancones basin. The district hosts a cluster of anomalously large deposits of base‐ and precious‐metal‐bearing volcanogenic massive sulphide (VMS) deposits around Tambogrande (Fig. 5.1, 5.2; Tegart et al., 2000; Chapter 1). On a grade‐
tonnage basis, the Tambogrande deposits are within the upper 3% of all VMS deposits of their type globally (Franklin et al., 2005) and are classified as bi‐modal mafic type according to Barrie and Hannington (1999). Analogous VMS deposits are of Archean age (e.g., Noranda camp, Gibson and Watkinson, 1990; Kidd Creek, Hannington and Barrie, 1999), Proterozoic age (e.g., Flin Flon, Syme and Bailes, 1993) and Phanerozoic age (e.g., Sibai and Gai, Herrington et al., 2005a,b). The deposits are the largest Mesozoic massive sulphide deposits in Perú and are the most potentially economic group of VMS deposits in South America. The Lancones basin contains three deposits (Fig. 5.2b): (i) TG1 with 109 Mt tons grading 1.6% Cu, 1.0% Zn, 0.5 g/t Au and 22 g/t Ag, (ii) TG3 with 82 Mt grading 1.0% Cu, 1.4% Zn, 0.8 g/t Au and 25 g/t Ag, and (iii) B5, with comparable massive sulphide intersections to TG1 and TG3, but defined tonnage or grade. The TG1 deposit also has an oxide zone of 16.7 Mt grading 3.5 g/t Au and 64 g/t Ag (indicated, inferred and probable mineral resources; Manhattan Minerals, 2002). The deposits remain undeveloped and, prior to this study, research on VMS mineralization and the volcanic host rocks in the region was limited (Injoque et al., 1979; Tegart et al. 2000; Winter et al., 2004). Presented in this paper are the first Pb, Sm‐Nd and Rb‐Sr isotopic data for Cretaceous marginal basin volcanic rocks in Perú. Isotope data from this study, along with published data for older metamorphosed igneous and sedimentary rocks, help establish a regional isotopic Chapter 5 Page 144 framework. The isotopic data illustrate the relative contributions of various continental crust and mantle sources in the petrogenesis of these volcanic rocks, which in turn assists in the construction and refinement of tectono‐magmatic models for the Cretaceous arc at the continental margin. Moreover, massive sulphide mineral (galena, pyrite) Pb isotope compositions from the VMS deposits establish lead (and by inference, other metal) sources for the Tambogrande VMS deposits which therefore aid in the development of metallogenic models. Empirical data from many VMS districts suggests a strong link between felsic volcanism and massive sulphide formation (Franklin et al., 2005), yet a range of ‘extensional arc’ tectonomagmatic scenarios and petrogenetic models have been proposed to explain this association (e.g., Lesher et al., 1986; Barrie et al., 1993; Lentz, 1998; Hart et al., 2004). This chapter examines the role of felsic and mafic volcanism in the genesis of VMS deposits at Tambogrande with respect to metal contribution and the hydrothermal convection system. The isotope data help test current VMS genetic models and assists in understanding the relationship between igneous petrogenesis of the volcanic suite and massive sulphide formation. 5.3 Regional Geology and Tectonic Setting The Tambogrande VMS deposits formed in the Lancones basin during the deposition of a latest Jurassic to Cretaceous volcano‐sedimentary sequence (Reyes and Caldas, 1987; Chap. 2.). The Lancones basin represents the northern portion of a much larger and segmented, Mesozoic continental margin rift basin extending southwards along the Peruvian coast (Huarmey‐Cañete Marginal Trough; Atherton et al., 1983; Benavides‐Cáceres, 1999) and also into southwest Ecuador (Jaillard et al., 1996). Similar arc‐related marginal rifts occur in Chile (Vergara et al., 1995) and Argentina (Dalziel, 1981; Hanson and Wilson, 1991). These marginal basins Chapter 5 Page 145 represent an episode of extensional tectonics in the first phase (~Mesozoic) of the Andean cycle in which Mariana‐type subduction permitted crustal attenuation, deposition of major marine sequences and eruption of large volumes of mafic‐dominated, subduction‐related volcanic rocks (Benavides‐Cáceres, 1999). The rift regime is temporally linked to the final opening of the South Atlantic in the Early Cretaceous and lasted until the Late Cretaceous when the geodynamical cycle shifted towards Andean‐type subduction in part due to active spreading in the South Atlantic. This resulted in the termination of marine sedimentation and the beginning of contractional tectonism, continental arc magmatism and a tectonic regime which continues to the present day. Located within the Huancabamba deflection, a major oroclinal bend in the Andes which separates the north‐northwest‐trending Peruvian Andes from the northeast‐trending Ecuadorian Andes, the Lancones basin records a sequence of accretionary events and contractional and extensional rotations (Mitouard et al., 1990; Fig 5.1). This tectonic development along the Peruvian‐Ecuadorian margin in Mesozoic times also involved shifts in convergence direction which were critical to the rifting event which formed the Lancones basin. In Jurassic times a southeast‐directed subduction zone was responsible for continental arc volcanism along the Ecuadorian segment (Litherland et al., 1994), whereas a sinistral transform system occurred along the Peruvian segment (Jaillard et al., 2000). During the Early Cretaceous a convergence shift to the northeast terminated the Ecuadorian magmatic arc and established the conditions for subduction along the Peruvian segment. Accretionary events during the Early Cretaceous also played a major role in the development of the margin, especially in northwestern Perú and Ecuador where allochthonous terranes are identified (Litherland et al., 1994). The Amotape terrane represents a Chapter 5 Page 146 microcontinental block comprised of early Mesozoic metamorphosed sedimentary, volcanic and plutonic rocks bordering the Lancones basin to the northwest and north (Mourier et al., 1988; Aspden et al., 1995; Litherland et al., 1994). This terrane includes meta‐granites with Middle to Late Triassic U‐Pb ages (Noble et al., 1997; this study, Appendix A) and in the northern segment in northernmost Perú and Ecuador, high‐pressure metamorphosed oceanic terranes comprised of mid‐ocean ridge (N‐MORB) and oceanic plateau basalts (e.g., Raspas metamorphic complex; Arculus et al., 1999; Bosch et al., 2002). Mourier et al. (1988) suggest the Amotape terrane arrived from the south and developed northeast‐trending dextral faults and clockwise rotation during accretion, the timing of which is constrained by K‐Ar cooling ages of ~132 to 110 Ma (Feininger and Silberman, 1982; Bosch et al., 2002). The accretionary event likely triggered the initiation of the new subduction zone outboard of the Amotape terrane resulting in the formation of the Lancones basin. Subsequent rifting which formed this basin probably utilized the suture between the Amotape Terrane and continental South America. To the southeast and east of the Lancones Basin (Fig. 5.2), the Paleozoic(?) Olmos Complex is probably a reactivated margin of the Amazonian craton (Macfarlane, 1999). This poorly defined terrane consists of pre‐Ordovician greenschist facies pelitic to psammitic rocks overlain by platform carbonate rocks of Triassic to Early Jurassic age. The Olmos Complex is considered to be equivalent to the Marañon Geanticline (Cobbing et al., 1981; Reyes and Caldas, 1987; Mourier et al., 1988; Litherland et al., 1994). The Olmos and Amotape Complexes represent the Jurassic to Early Cretaceous pre‐rift Andean margin and acted as topographical highs during Mesozoic times and controlled marine deposition in the Lancones Basin (Cobbing et al., 1981). Geophysical models constrain the crustal architecture of the continental margin of Perú and demonstrate a large arch‐like Chapter 5 Page 147 structure of dense material (3.0 g/cm3) that coincides with the Mesozoic volcanic rift sequences and separates the Amotape and Olmos continental blocks (Fig. 5.3; Couch et al., 1981; Jones, 1981). 5.4 Volcanic Stratigraphy of the Lancones Basin Rocks of the Lancones basin are exposed in northwestern Perú and southwestern Ecuador for more than 135 km along a northeast trend and approximately 150 km across the trend (Fig. 5.2). Tertiary cover blankets the basin in the southwest for an additional 50 kilometres. The basin can be subdivided into an eastern volcanic arc and western sedimentary forearc. The volcanic arc sequence, up to 80 km wide, is dominated by submarine mafic volcanic and volcaniclastic rocks, and transitions gradationally into forearc sedimentary rocks which dominate the western portion of the Lancones basin (Jaillard et al., 1999). The ~3 km thick Copa Sombrero Group (Chávez and Nuñez del Prado, 1991; Morris and Aleman, 1975; Jaillard et al., 1996, 1999) represents the western (forearc) turbiditic sub‐basin that temporally overlaps the volcanic arc sequence. Volcaniclastic rocks are intercalated with sedimentary rocks in the western part of the study area (upper part of the volcanic arc sequence) and suggest a transition to a forearc basin. The volcanic arc sequence in the Lancones basin comprises four formations with a total thickness of ~ 8 to 10 km (Figs. 5.4, 5.5; Chapter 2). These rocks include a wide spectrum of compositions and volcanic rock facies ranging from mafic to felsic volcanic rocks with lesser intermediate compositions, and also effusive lava flows to pyroclastic rocks with variable proportions of intercalated sedimentary rocks. In general, the sequence evolves from volcanic‐
dominated (lava flows) to volcaniclastic‐rich strata with a greater proportion of sedimentary interbeds. The sequence also appears to evolve from a deep to shallow marine and possibly Chapter 5 Page 148 subaerial facies based on an abundance of pyroclastic rocks, including felsic tuffaceous rocks, in the upper parts of the sequence. In Ecuador, the volcanic and volcaniclastic sequence has not been studied in detail and is described as a 2 to 3 km‐thick package of dominantly mafic pillow lavas and related volcaniclastic rocks (Jaillard et al., 1996). The volcanic arc sequence is subdivided into two main tectono‐volcanic phases based on depositional facies, composition and chronology, with the Cerro San Lorenzo Formation representing phase I and the Cerro El Ereo, La Bocana and Lancones formations defining phase II (Fig. 5.5; Chap. 2). The first phase is a mafic‐dominated sequence characterized by lava flows and associated breccias, with minor aphyric to weakly porphyritic felsic volcanic units, and is at least 2500 m thick. The rocks are interpreted to have been deposited in a deep water environment. U‐Pb zircon ages for volcanic rocks range from 99.1 to 104.8 Ma (Chap. 2). The phase I volcanic sequence is also of economic interest as it hosts all VMS deposits in the Lancones basin. The phase II volcanic cycle is an 8 km‐thick sequence of mafic to felsic volcanic and volcaniclastic rocks with subordinate calcareous and siliciclastic sedimentary rocks. These rocks were deposited in a relatively shallow water setting. U‐Pb zircon ages range from 99.3 to 91.1 Ma for volcanic rocks of Phase II (Chap. 2). The Lancones Formation, the youngest rocks of the phase II cycle, includes abundant volcaniclastic (epiclastic) rocks at its base and grades upwards into dominantly sedimentary rocks marking a transition to the forearc domain. The volcanic arc sequence shows no effects of dynamic metamorphism, which ranges from zeolite to lower greenschist facies, and primary textures are generally well preserved. Chapter 5 Page 149 5.5 Andean Isotopic Framework Metallogenic terrane characterization of the Central Andes defines three geologic provinces with distinctive sulphide Pb isotope signatures that are attributed to variable geology, magmatism and metallogeny (Macfarlane et al., 1990; Macfarlane, 1999). Province I is defined as the coastal region and includes mineral deposits associated with Jurassic‐early Tertiary volcanic arc rocks and the Coastal batholith or Perú and Chile (Fig. 5.6). Pb isotope compositions of province I plot at or below Stacey and Kramers (1975) average crustal growth curve and define a gently sloping trend with moderate range of 206Pb/204Pb values (Fig. 5.7). Province III overlaps the Eastern Cordillera and Altiplano and comprises Paleozoic sedimentary rocks, underlying Proterozoic(?) basement, and igneous rocks of dominantly crustal derivation, which record events along the Gondwanan margin. Pb isotopic compositions of province III define a broad array of radiogenic Pb isotope ratios that plot mostly above Stacey and Kramers (1975) average crustal growth curve and extend over a large range of 206Pb/204Pb values. Province II, broadly defined as the ‘high Andes’ of Perú, includes ore deposits that occur between provinces I and III and which are associated with Oligocene and younger volcanic rocks. Pb isotope ratios in province II yield a narrow range of 206Pb/204Pb and form steep Pb isotope arrays which approximate a mixing trend. The Pb isotope reservoirs responsible for this mixing trend are inferred to be a homogenized mantle‐dominated province I‐type source and a radiogenic, crustal‐dominated province III‐type source (Macfarlane et al., 1990; Macfarlane, 1999). More subtle variations in Pb isotopic compositions are reflected as latitudinal variations within the provinces. For example, the subprovince Ib in central Perú displays higher 206
Pb/204Pb than either of the subprovinces Ia and Ic which extend further south. Likewise, Chapter 5 Page 150 subprovince IIIb lies farther north than subprovince IIIa, and yields significantly higher 206
Pb/204Pb ratios. Variations in 207Pb/204Pb and 208Pb/204Pb are less pronounced within a particular Pb isotopic province. Sedimentary and basement rocks of northern Perú yield higher 206
Pb/204Pb ratios than basement rocks in southern Perú and account, at least in part, for the diverse variation in sulphide Pb compositions in the region (Macfarlane, 1999). In northwestern Perú and southwestern Ecuador, distinctive Pb isotopic signatures also characterize major geologic units that are integral components of the accretionary and magmatic arc history of the Lancones basin. A schematic tectonic model of the Cretaceous northwestern Perú region and the spatial distribution of isotope samples is shown in Figure 5.8. The Olmos Complex basement rocks yield a broad range of Pb isotope ratios which are also the most radiogenic of all geological units in the region and are incorporated in subprovince IIIb (Fig. 5.9). Likewise, the Cretaceous platform sedimentary rocks, which overly the Olmos Complex and were likely derived from it, display similar radiogenic Pb compositions though with steeper Pb isotope arrays, which mostly mimic subprovince IIIb and possibly province II. A single sample from an Amotape Complex meta‐granite (this study) yields Pb isotopic ratios that are similar to the least radiogenic of the Olmos Complex. Within the Amotape accretionary terrane, Jurassic‐Cretaceous meta‐basalts from the Raspas metamorphic complex in southwestern Ecuador yield mid‐ocean ridge basalt (MORB) Pb isotopic compositions that are similar to Pb values in present day MORB in the east Pacific Ocean. Meta‐sedimentary rocks from the Raspas metamorphic complex, however, yield relatively radiogenic Pb compositions that overlap with the lowermost ratios of the Cretaceous platform sedimentary rocks. Mafic‐
intermediate plutons of the Coastal batholith of the northern Andes plot just below Stacey and Kramers (1975) average crustal growth curve, whereas felsic plutons are relatively more Chapter 5 Page 151 radiogenic with isotope ratios greater than average crust and overlap the least radiogenic compositions of the Olmos Complex. No Pb isotope data are reported in the literature for the volcanic rocks or sulphide deposits of the Lancones Basin. The Sr and Nd isotopic framework for the Andes is less well established in comparison to Pb, however, data are available for major geologic units in the northern Andes (Fig. 5.10A). The Olmos Complex, basement rocks to the Cretaceous marginal rift basins, yield relatively high 87
Sr/86Srpresent‐day (0.732‐0.741) and low εNd present‐day (‐12.1 ‐ ‐11.5) typical of highly radiogenic lower crustal rocks, whereas overlying Cretaceous sedimentary rocks have comparatively lower 87
Sr/86Srpresent‐day (0.708‐0.725) but similar εNd present‐day (‐16.5 ‐ ‐11.7; Macfarlane, 1999). In the Jurassic‐Cretaceous Raspas Metamorphic Complex, metapelite yields moderately high initial 87
Sr/86Srpresent‐day (0.716‐0.718) and εNd present‐day (‐6.8 ‐ ‐9.9), whereas metabasalts yield MORB‐
like 87Sr/86Sr (0.707) and εNdpresent‐day = 10.6 (Bosch et al., 2002). Post‐rift granitoid intrusions (87Sr/86Sr = 0.70351 ‐ 0.7051) are generally considered to be the result of subduction zone enrichment of a subcontinental mantle source and show limited contamination with radiogenic upper crustal rocks (Soler and Rotach‐Toulhoat, 1990). Sr and Nd data are not available for Cretaceous volcanic rocks from the Peruvian marginal rift basins. 5.6 Pb, Sm‐Nd and Rb‐Sr Isotope Geochemistry Twenty‐two Pb isotopic measurements were made on sulphide minerals separates from the TG1, TG3, and B5 VMS deposits (n=6) and whole rock samples of syn‐mineralization volcanic rocks in the vicinity of TG1 and TG3 (n=10) from the Cerro San Lorenzo Formation. Pb isotope data are also reported from post‐mineralization basalt samples from the Cerro El Ereo Formation (n=2), feldspar separates from Amotape metagranite basement rock (n=1) and post‐
mineralization Cretaceous granitic rocks (n=3). Additional isotopic constraints are provided by Chapter 5 Page 152 12 Sm‐Nd and Rb‐Sr isotope analyses of the volcanic rocks. Sample locations are shown on the map in Figure 5.4 and descriptions are provided in Tables 5.1 and 5.2. All samples were analyzed at the Pacific Centre for Isotopic and Geochemical Research (PCIGR), Department of Earth, Ocean, and Atmospheric Sciences, University of British Columbia, Vancouver, B.C., Canada. 5.7 Analytical Methods 5.7.1 Pb Isotope Analysis, Mineral Separates 1) Samples of sulphide separates were prepared from 10‐50 mg of hand‐picked pyrite crystals, and were leached in dilute nitric acid and then hydrochloric acid to remove surface contamination before dissolution in dilute nitric acid. 2) Samples of feldspar separates were prepared from 10‐50 mg of hand‐picked feldspar crystals, and these were leached in dilute hydrochloric acid and then dilute hydrofluoric/hydrobromic acids to remove surface contamination prior to dissolution in hydrofluoric acid. Separation and purification of Pb employed ion exchange column techniques. The dissolved samples were evaporated and heated to dryness, converted to chloride, then dissolved in dilute hydrobromic acid. This solution was passed through ion exchange columns, and the lead eluted in 6N hydrochloric acid. 3) Small (≤1 mm), clean galena crystals were dissolved in 2N hydrochloric acid, then rinsed in 4N hydrochloric acid to clean the resulting lead chloride crystals. Approximately 10‐25 ng of the lead in chloride form was loaded on a rhenium filament using a phosphoric acid‐silica gel emitter, and isotopic compositions were determined in peak‐switching mode using a modified VG54R thermal ionization mass spectrometer. The measured ratios were corrected for instrumental mass fractionation of 0.12%/amu (Faraday collector) per mass unit based on repeated measurements of the N.B.S. SRM 981 Standard Isotopic Reference Material and the Chapter 5 Page 153 values recommended by Thirlwall (2000). Errors were numerically propagated including all mass fractionation and analytical errors, using the technique of Roddick (1987). All errors are quoted at the 2σ level. The analyses have been plotted with the average crust model lead growth curve of Stacey and Kramers (1975) for comparison. Age assignments follow the time scale of Harland et al. (1990). The total procedural blank on the trace lead chemistry was 64 pg. 5.7.2 Pb, Rb‐Sr, Sm‐Nd Isotope Analysis, Whole Rock Samples Isotopic composition measurements were determined on a Finnigan Triton thermo‐
ionization mass spectrometer (TIMS; Sr, Nd) and on a Nu Instruments (Nu 021) multiple collector inductively coupled plasma mass spectrometer (MC‐ICP‐MS; Pb) at the Pacific Centre for Isotopic and Geochemical Research (PCIGR) at the University of British Columbia. The complete methodology is described by Weiss et al. (2005). Sr and Nd compositions were measured in static mode multicollection with relay matrix rotation (the “virtual amplifier” of Finnigan) on a single Ta and double Re‐Ta filament, respectively. The data were corrected for mass fractionation using 86Sr/88Sr = 0.1194 and 146
Nd/144Nd = 0.7219, respectively. Fifty‐five analyses of the NIST SRM 987 Sr standard and seventy‐six analyses of the La Jolla Nd standard made during the course of this study have mean values of 87Sr/86Sr = 0.710250 ± 12 (2SD) and 143Nd/144Nd = 0.511853 ± 16 (2SD), respectively. A single analysis typically consists of 135 cycles (9 blocks of 15) to allow a full rotation of the virtual amplifier. Pb isotopic compositions were analyzed by static multicollection. The collector array on the Nu Plasma is fixed and a zoom lens is employed to position the masses in the collectors. The central collectors (H4‐L2) are 1 amu apart while the outer collectors (H6, H5, L3, L4 and L5) are 2 amu apart. Masses 208 to 202 are measured in collectors H4 to L2. Chapter 5 Page 154 5.8 Results Initial 87Sr/86Sr (87Sr/86Srί) and initial 143Nd/144Nd (143Nd/144Ndί) isotopic ratios, as well as initial εNd (εNdί) for whole rock samples have been calculated to 100 Ma, the approximate age of the volcanic rocks in this study (Chapter 2). 143Nd/144Ndί and εNdί were calculated using the present day values of 143Nd/144Nd = 0.5123638 for the chondrite‐uniform reservoir (CHUR) (Hamilton et al., 1983), 147Sm/144Nd = (Sm/Nd)*0.602 and a decay constant λ of 6.54 x 10‐12yr‐1 (Faure and Mensing, 2005). 87Sr/86Srί was determined using a 87Rb/86Sr ratio = (Rb/Sr)*2.8935 (Faure and Mensing, 2005) and a decay constant λ of 1.42 x 10‐11yr‐1. 208,207,206Pb/204Pb values are given as present day measured values as precise Pb and U concentration data are not available for these samples. Sr, Nd, and Pb isotope data are listed in tables 5.2 and 5.3 5.8.1 Volcanic Rocks Mafic and felsic volcanic rocks yield distinctively different Pb isotope compositions that do not overlap on 207Pb/204Pb vs 206Pb/204Pb and 208Pb/204Pb vs 206Pb/204Pb plots (Figs. 5.7 & 5.9). Moreover, these rocks define markedly distinctive arrays with different slopes. Basalt samples define a line with a steep slope and yield relatively less radiogenic Pb ratios which overlap the lower 206,207,208Pb/204Pb section of Province IIIb field. In contrast, felsic volcanic rocks plot along an array with a relatively flat slope, yield a broad range of 206Pb/204Pb values, and have highly radiogenic Pb ratios which correspond to the higher Pb ratios of the Province IIIb field. Despite being erupted in the same setting at approximately the same time, the isotopic variations suggest variable processes in magma genesis for these units due to the varying influence of relatively old and radiogenic continental crust. Pb isotope ratios for three whole‐rock basalt samples from the Cerro San Lorenzo Formation plot above Stacey and Kramers (1975) crustal growth curve and yield a narrow range of Pb isotope compositions (206Pb/204Pb = 18.94 – 19.02). Albeit only a small number of Chapter 5 Page 155 samples, the group forms a steep array (slope = 0.676), though a fourth basalt sample is an outlier with higher 206Pb/204Pb (19.36) and is intermediate between the main basalt cluster and the felsic volcanic rocks in Pb isotope space. This outlier also has slightly higher Th (i.e., 1.04 ppm versus 0.25‐0.89 ppm Th in the other basalt samples). This sample is from hanging wall rocks at TG3 and therefore may have been contaminated by felsic volcanic strata which underlie this unit. Pb isotope ratios of whole‐rock basalt samples from the stratigraphically higher Cerro El Ereo Formation are similar to the main cluster of basalts from the Cerro San Lorenzo Formation. The basalt samples have Pb isotope compositions similar to continental crust of the Olmos Complex and overlying Cretaceous platform sedimentary rocks. The basalt samples also have equal or slightly greater 207Pb/204Pb ratios than felsic phases of the Coastal batholith which intrude them, suggesting the basalts were generated in a slightly higher U/Pb environment than subsequent granites in the same region. Compared to the Pb provinces of Macfarlane et al. (1990), these rocks have Pb isotope ratios that overlap province IIIb, though they also have similar Pb isotope to province II, except for slightly greater 206Pb/204Pb values. Six whole‐rock felsic volcanic samples associated with VMS deposits at TG1 and TG3, which include syn‐mineralization rhyolite lava flows, post‐mineralization rhyolite dykes, and post‐
mineralization dacite lava flows, yield much higher 206Pb/204Pb (19.64 – 20.35) and 206Pb/204Pb (39.40‐39.79) than contemporaneous basalt. The data define a line with a relatively flat slope (0.056, correlation coefficient = 0.98) that projects above the average crustal growth curve of Stacey and Kramers (1975) and yields an errorchron age of 452 Ma (Fig. 5.9B; calculated using Isoplot 3.0; Ludwig, 2003). This errorchron does not have any geological significance. These samples have Pb isotope ratios comparable to the Olmos (and Marañón) basement rocks and overlying platform sedimentary rocks, and overlap the subprovince IIIb Pb fields. Chapter 5 Page 156 Rb‐Sr isotope data define a relatively narrow range of whole rock 87Sr/86Srί values ranging from 0.703675 to 705860 (Fig. 5.10). Felsic volcanic rocks have the most restricted range in values, with 87Sr/86Srί ranging from 0.704374 to 0.705063. A coherent array defined by the felsic volcanic rocks yields a crude errorchron of ~65 Ma (calculated using Isoplot 3.0; Ludwig, 2003) which does not have any geologic significance (Fig 5.10B). Basalt from the Cerro San Lorenzo Formation in the footwall to massive sulphide deposits has been hydrothermally altered by the mineralizing fluid and has depleted Sr concentrations (111‐150 ppm) in comparison to relatively unaltered hanging wall basalt (350‐483 ppm). Sr contents in basalt correlate negatively with loss on ignition values (L.O.I.). Moreover, as a result of hydrothermal alteration, Sr mobility, and Sr isotope exchange, footwall basalt has higher initial 87Sr/86Srί values (0.705208‐0.705235) than hanging wall basalt (0.703675‐0.704276; Fig 5.10A). These ‘modified’ higher values are slightly greater than the felsic volcanic rock Sr isotope compositions. Basalt from the hanging wall, which is a better estimate of the primary igneous Sr isotope composition, has Sr isotope ratios that are somewhat lower than the group of felsic volcanic rocks. Younger basalt from the Cerro El Ereo Formation has Sr isotope compositions that are similar to basalt from the Cerro San Lorenzo Formation. Sm‐Nd isotope data yield positive εNdί values for all samples, though there is a systematic variation between volcanic rocks of basaltic to rhyolitic compositions. Cerro San Lorenzo Formation basalt has relatively low Sm (1.22 – 2.61 ppm) and Nd (3.26 – 9.52 ppm) contents and yields a narrow range of εNdί (4.73 ‐ 5.93). Similarly, Cerro El Ereo Formation basalt also has relatively low Sm (1.17 – 2.38 ppm) and Nd (3.38 – 7.23 ppm) contents and a slightly broader range of εNdί (3.82 – 6.83). In contrast, felsic volcanic rocks have higher Sm and Nd contents and somewhat lower εNdί values. Syn‐mineralization rhyolite has the highest Sm and Nd Chapter 5 Page 157 contents (5.76 ‐ 6.72 ppm and 18.72 ‐ 22.93 ppm, respectively) and εNdί values (4.40 ‐ 5.59) that overlap with the mafic volcanic rocks. However, both post‐mineralization dacite (Sm = 4.30 ‐ 4.32, Nd = 16.08 – 16.62, εNdί = 3.01) and rhyolite dykes (Sm = 4.01 – 5.28, Nd = 16.85 – 21.45, εNdί =1.85 ‐ 1.95) yield distinctly more evolved Nd isotopic signatures. In summary, the volcanic rocks display a broad range of positive εNdί values (1.85 to 5.93) that are beyond the range of analytical error (±0.8) and are evidence of depleted mantle‐
derived magma sources with variable influence of more radiogenic crustal sources. This is consistent with Sr isotope systematics whereby the volcanic rocks have Sr isotopic compositions close to bulk‐earth. In Sr‐Nd isotope space all volcanic rocks from this study lie above the mantle array and are more enriched than Mesozoic or present‐day Pacific MORB, trending towards the highly enriched fields for crustal rocks (Fig. 5.11A). The lead isotope systematics reflect the crustal component more effectively as all volcanic rocks in the study are characterized by high Pb isotope ratios. 5.8.2 Massive Sulphide Deposits Six galena and pyrite mineral separates from the TG1, TG3 and B5 massive sulphide deposits yield a narrow range of Pb isotope ratios (206Pb/204Pb = 18.87‐18.93) that partly overlap when analytical uncertainty is accounted for. A steep linear array (slope = 0.676, correlation coefficient = 0.98) plots within the Pb isotope field for Olmos basement rocks and is also close to the field of Cretaceous sedimentary rocks (Fig. 5.9B). The sulphide Pb isotope ratios are similar to, but slightly greater, than those of the samples from the Cerro San Lorenzo Formation basalt, their host rocks. The sulphide separates yield unreasonably old model ages and no chronologic significance is attached, as the trend line does not intersect the average crustal growth curve of Stacey and Kramers (1975). Rather, the trend reflects either mixed Pb sources Chapter 5 Page 158 or analytical uncertainty due to mass fractionation. Steep trends are typical of province II Pb, which are interpreted to be mixed sources (Mcfarlane et al., 1990). 5.9 Discussion 5.9.1 Volcanic Rock Petrogenesis Pb, Rb‐Sr and Sm‐Nd isotope systematics can be used to elucidate the variability of the magma sources for Cretaceous volcanic rocks in the vicinity of VMS deposits in the Lancones Basin and suggest a depleted upper mantle (i.e., mantle‐wedge) source for volcanic rocks which was variably hybridized by crustal assimilation of older continental rocks. The crustal isotopic signature is most notable in the Pb isotopic data, but is also evident in the Sr and Nd isotope data, and in all systems (i.e., Pb, Sr, Nd isotopic), the felsic volcanic rocks yield the most radiogenic signatures. Mafic volcanic rocks in the Lancones basin display geochemical characteristics of weakly calc‐alkaline to weakly tholeiitic oceanic arc volcanic rocks related to subduction zone enrichment of an N‐MORB mantle source (Chap. 4). Isotopically, the mafic rocks are typical of arc volcanic sequences and are displaced from MORB signatures to more enriched isotopic compositions. There are two possible explanations for the isotopic enrichment which caused the basalt to inherit the signature of crustal rocks. One possibility is via the subduction and recycling of sedimentary rocks derived from ancient cratonal sources. Mixing of depleted mantle source and slab‐derived component is well established in oceanic arc volcanic rocks (McCulloch and Gamble, 1991; Hawkesworth et al., 1993). The assimilation of crustal rocks, either metamorphic basement or platform sedimentary rocks, by the arc magma on ascent into the upper crust is a viable alternative. Since both Jurassic‐Cretaceous metasedimentary rocks of the Raspas metamorphic complex, which are equivalent to the subducted sedimentary rocks, Chapter 5 Page 159 and the Olmos metamorphic complex have overlapping Pb isotope ratios, either or both are possible sources for the isotopic enrichment of the mafic rocks. Felsic volcanic rock yield Pb isotope compositions are significantly more radiogenic than the basalt samples and also show significant variation in 206Pb/204Pb ratios and linear arrays on the Pb isotope diagrams. These data do not plot along a growth curve nor do they define an isochron; rather the relatively linear trend of the data indicates a mixing line between two sources. The lower endmember of this mixing line intersects the basalt field which is consistent with geochemical affinities signifying mantle derivation (i.e., M‐type magmas; Chap. 4), either as modified products of basalt‐parent magmas or as partial melting of juvenile arc (mafic) crust during underplating of the mid‐upper crust. The upper endmember of the trend is within the field of Olmos Complex basement rocks and represents homogenized Pb from this relatively old crust. Since there are no other known reservoirs with Pb signatures as radiogenic as the Olmos basement, and considering the proximity to this crustal material, the Olmos is the most likely radiogenic component for the felsic lavas. That mafic and felsic lavas, erupted contemporaneously within the same volcanic edifice, have such variable isotopic signatures suggests that the parent melts were generated or hybridized at different crustal levels. The effect of crustal contamination is further illustrated geochemically in plots of εNdί versus Th/Yb and 206Pb/204Pb and versus Th/Yb (Fig. 5.12A,B). Higher Th/Yb reflects the increased effect of crustal contamination which in this dataset shows a strong negative correlation with εNd values. There is no correlation between Th/Yb and 206
Pb/204Pb, though all felsic rocks yield higher Pb ratios when compared with basalts. Despite higher Pb isotope ratios for the syn‐mineralization rhyolite, Th/Yb and εNd values broadly overlap with the basalt data. In contrast, post‐mineralization dacite flows and rhyolite dykes Chapter 5 Page 160 are clearly distinguished by lower εNd values and higher Th/Yb due to greater crustal assimilation in the parent magmas. Further direct evidence for crustal contamination includes minor inheritance of older zircons in some of the volcanic rocks (Chapter 2). Granites from the Coastal Batholith are much less radiogenic than older felsic volcanic rocks but are isotopically similar to basalt. The granitoids of the coastal region (Western Cordillera) are voluminous and yield isotopic compositions that suggest there was little influence from radiogenic old crust (Soler and Rotach‐Toulhoat, 1990). Further, Eastern Cordillera batholiths contain little evidence from contamination of old continental crust, despite traversing basement rocks up to >50 km thick (Petford et al., 1996). The isotopic data presented here have implications for the genesis of magma and the processes of crustal growth along the Peruvian margin. Felsic volcanic rocks in the Cretaceous rift basins were generated in part by small batch crustal recycling in the upper crustal sectors and are not voluminous mantle‐
derived, evolved melts such as the subsequent granitic rocks. 5.9.2 Metal Sources for VMS Deposits Pb isotope data from different deposits within individual VMS camps commonly plot as tight clusters on Pb isotope plots (Thorpe, 1999). At Tambogrande, Pb isotope data from the three massive sulphide deposits are broadly similar with many of the values being identical, within analytical error. This suggests a large hydrothermal system with homogenized Pb isotopic composition was responsible for the formation of the three VMS deposits. Such as interpretation is consistent with the equivalent stratigraphic position and uniform features for these deposits (Chapter 3). Using Pb isotope systematics, an approximation of the relative metal contributions from the host rocks to the sulphide deposits can be determined. Chapter 5 Page 161 As described in the previous section, mafic volcanic rocks form a cluster on the Pb isotope plots and are the least radiogenic of the volcanic rocks (Fig. 5.13). Felsic volcanic rocks, however, are highly radiogenic and plot on an array that indicates mixing between a basalt source and radiogenic crustal rocks. Massive sulphide Pb isotope data plot between the mafic and felsic data, though the values are closest to the basalt suite (using conceptual time‐
adjusted 206Pb/204Pb values; 100 Ma), based on Stacey and Kramers (1975) average crustal growth curve with a µ value of 9.85). Assuming the steep linear trend of the sulphide Pb isotope array is not solely a function of analytical error, two main components are inferred. The upper endmember is probably homogenized Pb from crustal material, such as the Olmos Complex or sedimentary rocks overlying the metamorphic basement. That the most radiogenic sulphide samples plot near the upper limits of the Olmos Complex field (Fig. 5.9) suggests that materials in the crustal column may be more radiogenic than previously recognized. Alternatively, as noted by Macfarlane (1999), leachates of Olmos Complex rocks yield more radiogenic 206Pb/204Pb (also for 208Pb/204Pb, less so for 207Pb/204Pb) than whole‐rock samples and therefore hydrothermal leaching may have scavenged Pb with higher average Pb ratios. The lower component of the sulphide mineral Pb array is similar to the basalt Pb isotope field but is displaced to slightly higher 206Pb/204Pb ratios, specifically in the direction towards felsic volcanic rocks (Fig. 5.13). Although the volcanic strata are largely dominated by basalt, the slight displacement to higher 206Pb/204Pb ratios suggests the volcanic component reflects mineralizing hydrothermal fluid homogenization of the Pb isotopic compositions of both mafic and felsic volcanic rocks which comprise the footwall. The homogenized volcanic rock Pb isotope composition is determined by the intersection of the steep sulphide mineral Pb isotope array with the relatively flat felsic volcanic rock array (essentially the same as the mixing line Chapter 5 Page 162 between syn‐mineralization rhyolite and basalt). The intersection of these lines represents the approximate homogenized Pb composition of the footwall volcanic rocks (206Pb/204Pb ≈ 18.90, 207
Pb/204Pb ≈ 15.67, labeled in Fig. 5.13). Using approximate average Pb isotopic compositions for the mafic and felsic volcanic rocks, as well their respective Pb contents, it is possible to estimate the relative Pb contributions of each of these volcanic rock units to the VMS deposits. Based on average 206Pb/204Pb ratios for basalt = 18.80, rhyolite = 19.95, and massive sulphide minerals = 18.90, as well as an approximate Pb content for rhyolite that is nominally 2.5 times greater than basalt (rhyolite ≈ 5 ppm, basalt < 2 ppm), massive sulphide Pb contributions are 3% for rhyolite and 97% for basalt. This is consistent with field observations which suggest an almost wholly mafic footwall sequence in the Cerro San Lorenzo Formation. Therefore, despite an intimate spatial and temporal association, contemporaneous felsic volcanic rocks apparently provided a negligible contribution to the Pb metal budget of the VMS deposits at Tambogrande. The proportion of other metals (copper, zinc, gold) from the mafic volcanic rocks is perhaps even higher considering the abundance of these metals in the mafic rocks in comparison to the felsic rocks. The data suggest that the bulk of metals for the Tambogrande VMS deposits were derived via hydrothermal leaching of the underlying mafic volcanic strata and/or from a magmatic volatile phase from mafic subvolcanic intrusions of the same isotopic character. 5.9.2 Possible Linkages Between Petrogenesis of the Felsic Volcanic Suite and VMS Formation Current VMS genetic models suggest a link between the nature and extent of VMS hydrothermal circulation systems and contemporaneous igneous rocks (Lydon, 1988; Franklin, 1995; Franklin et al., 2005; Galley, 1993). All VMS genetic models require a thermal source to drive the hydrothermal convective system (and possibly contribute metals), that is generally Chapter 5 Page 163 considered to be new magma input to the upper crust. Heat‐ and fluid‐flow modeling indicates that substantial thermal input into the upper crust is required in order to sustain a hydrothermal convection system capable of producing giant massive sulphide deposits (Cathles et al., 1997; Barrie et al., 1999). Therefore, the thermal input required is most likely accommodated by mantle‐derived intrusions (ultramafic‐mafic) emplaced within the upper crust. These intrusions are responsible for the generation of the hydrothermal system, but also generate rhyolitic magmas through partial melting of crust above the intrusion (Fig. 5.14). Felsic volcanic rocks at Tambogrande have geochemical signatures that are characterized by relatively enriched HREE and Y and weak LREE/HREE fractionation (Chap. 4) and are considered to be high temperature melts (Chap. 2), similar to felsic volcanic rocks from other VMS camps globally (Lesher et al., 1986; Barrie et al., 1993; Barrett and Maclean, 1994; Lentz, 1998). These geochemical compositions are consistent with partial melting of hydrated mafic crust at relatively high temperature and low pressure conditions in equilibrium with hornblende (or clinopyroxene at higher temperatures) and plagioclase (Hart et al., 2004, and references therein). Using this model, Pb isotope data of the massive sulphide deposits and volcanic rocks at Tambogrande are used to further constrain the relationship between VMS formation and the petrogenesis of the associated igneous rocks. Any proposed genetic model must incorporate the fact that felsic volcanic rocks are not simply the result of fractionation from a basaltic parental magma, nor are they solely melts of continental crust. Rather, the felsic rocks are a two‐component hybridized melt comprised of (i) enriched mantle (~arc basalt source) and (ii) old continental crust, as evidenced by the radiogenic Pb isotope ratios. The mantle component may have been either a melt of the same parentage as the mafic lavas or a melt derived from Chapter 5 Page 164 partial melting of juvenile arc crust and therefore of the same geochemical affinity. In either case, a mafic source melt was able to assimilate basement rocks at mid‐upper crustal levels, and this implies a significant thermal input to the crust at relatively shallow levels. As outlined above, Pb contribution from the felsic volcanic rocks was trivial, most likely because felsic strata are volumetrically insignificant within the footwall sequence. The Pb isotope data also imply that negligible metal was contributed from the basement rocks which were responsible for the isotopic signature of the felsic volcanic rocks. The maximum depth reached by a hydrothermal system in oceanic crust is determined in part by the maximum depth of brittle fracture permeability and has been estimated to be 8‐10 km in oceanic crust (Barrie et al., 1999). If the source region of partial melting and/or crustal contamination producing the rhyolitic melts was below the maximum depth of the hydrothermal system, mineralizing fluids would not have encountered the regions where continental basement material occurs. Indeed, the hydrothermal system was limited to upper sections consisting of juvenile arc crust of basaltic composition as evidenced by the basalt‐dominant Pb isotope signature of the massive sulphide minerals. The close spatial and temporal association of felsic rocks with penecontemporaneous VMS deposits at Tambogrande is symptomatic of the thermal anomaly present at the time of massive sulphide formation. High‐level crustal input of subduction‐related mafic magma, either directly or via partial melting of juvenile arc crust, assimilated continental crust to generate isotopically distinct felsic melts. These mafic melts also fuelled the hydrothermal system responsible for the VMS deposits by underplating the upper crust and perhaps also by forming nested intrusions at even higher levels (Fig. 5.13). Felsic magmas ascended rapidly from the area of generation in the upper or mid‐crust(?) and were erupted in the general vicinity of Chapter 5 Page 165 hydrothermal venting on the seafloor and VMS formation. Felsic magmas likely utilized the same pathways that enabled focused hydrothermal fluid flow and VMS formation. The felsic volcanic rocks are therefore considered to have played a passive role in the formation of VMS deposits at Tambogrande and are only indirectly genetically linked. 5.10 Summary and Conclusions The Tambogrande VMS deposits are associated with volcanic rocks of variable radiogenic (Pb, Sr, Nd) isotopic compositions that indicate two major components: (i) a depleted upper mantle, MORB‐type source (i.e., the mantle‐wedge) and (ii) old continental crust. Basalt from the Lancones Basin displays Pb isotopic compositions that plot above the average crustal growth curve and suggest mixing between a depleted mantle source and continental crust, yielding an enriched‐mantle signature. This radiogenic Pb component was likely added through assimilation of sedimentary rocks in the subduction zone and possibly also through contamination during ascent through the crust. Felsic volcanic rocks have more radiogenic Pb, Sr and Nd isotope ratios than basalt and define a mixing line between the basalt and a highly radiogenic source. The crustal signature is most evident in the Pb isotope systematics but is also present in the Sr and Nd data. The basalt component is either a primary melt equivalent to the basalt or re‐melted juvenile arc (mafic) crust, whereas the radiogenic source is old continental crust such as the Olmos Complex or sedimentary rocks of similar provenance. Hybridization of the mafic magmas resulted from the assimilation of continental crust in the mid‐upper crustal column. That these bimodal volcanic rocks were erupted essentially in the same space and time attests to the complexity of VMS‐
associated volcanic magmas. This interpretation is consistent with the inferred geodynamic setting of a continental margin volcanic arc‐rift. Chapter 5 Page 166 The uniformity of massive sulphide Pb isotope compositions from the TG1, TG3 and B5 deposits suggests homogenization of Pb in the VMS‐forming hydrothermal system. The deposits either formed from a common hydrothermal system or derived metals from the same source rocks. Based on the Pb isotope data, metal sources for the VMS deposits are dominated by basaltic footwall rocks, whereas felsic volcanic rocks contributed insignificant amounts of metal. This is consistent with observations of a footwall sequence composed nearly entirely of mafic volcanic rocks. The Pb isotope data occur along a steep array on Pb isotope plots, indicating two‐component mixing, very similar to the Province II Pb isotope field of Mcfarlane et al. (1990). The least radiogenic endmember is the mafic‐dominant volcanic sequence, whereas the more radiogenic endmember is homogenized lead from the continental crust. Genetic models for VMS formation at Tambogrande favour hydrothermal leaching of the footwall mafic volcanic rocks (and/or a metal contribution from magmatic devolatilization of a mafic intrusion). The close spatial and temporal association of felsic volcanism with the VMS deposits is linked to the thermal anomaly generated at mid‐upper crustal levels by the introduction of mafic magmas which likely underplated the upper crust. These magmas, or partial melts of the existing juvenile arc crust, were subsequently hybridized by older continental crust and resulted in high‐temperature felsic magmatism. The heat input into the upper crust was sufficient to generate large‐scale hydrothermal convective circulation that resulted in deposition of massive sulphides at and on the seafloor. Therefore, the petrogenesis of the felsic volcanic rocks associated with the Tambogrande VMS deposits has no causal relationship with these deposits, but it is a symptom of anomalous heat input to the upper crust during VMS deposit formation. This passive role of felsic volcanism associated with VMS Chapter 5 Page 167 systems is likely a common feature to many other VMS camps, especially in the mafic‐
dominated class. Chapter 5 Page 168 Figure 5.1. Morphostructural units of the Peruvian Andes (modified after Benavides‐Cáceres, 1999). Cretaceous marginal basins ‐ Lancones (LB), Huarmey (HB) and Cañete (CB) basins ‐ are superimposed. Also shown are the locations of VMS deposits and prospects (circles) (data from Steinmüller et al., 2000). Chapter 5 Page 169 Figure 5.2. A. Location map for the Tambogrande project B. Regional map showing major tectonostratigraphic units of coastal northwestern Perú. The locations of VMS deposits (TG1, TG3, and B5) in the Tambogrande area are also shown and field area of this study outlined (see Fig. 5.3 for a detailed map). Modified after Jaillard et al. (1999), Tegart et al. (2000). Chapter 5 Page 170 Figure 5.3. Location map and simplified cross sections along the Peruvian continental margin based on gravity modeling and seismic data from Couch et al. (1981) and Jones (1981). Chapter 5 Page 171 Figure 5.4 (next page). Regional geologic map for the Tambogrande area of the Lancones Basin. The location of VMS deposits TG1, TG3, and B5, where the bulk of the isotope samples were collected are shown Other individual samples from the region are labeled. Map projection is WGS 84 (World Geodetic System), UTM Zone 17 Southern Hemisphere. Chapter 5 Page 172 Chapter 5 Page 173 Figure 5.5. Schematic stratigraphic column of the volcanic arc sequence of the Lancones basin. Inset section shows a more detailed schematic section of the VMS‐bearing sequence at Tambogrande. Chapter 5 Page 174 Figure 5.6. Map depicting Pb isotope provinces of the Andes (modified after Macfarlane et al., 1990; Tosdal et al, 1999). The location of Tambogrande and other VMS deposits within the Cretaceous marginal basins of Perú are shown. Chapter 5 Page 175 Figure 5.7. Thorogenic (A) and Uranogenic (B) Pb isotope diagrams for data from the Lancones basin and fields for Pb isotope provinces of the Andes (after Macfarlane et al., 1990). All data points are from this study. Symbols for rocks and sulphide samples and fields for Pb provinces are given in inset boxes in B. S & K = Stacey and Kramers (1975) growth curve. Chapter 5 Page 176 Figure 5.8. Schematic east‐west cross section through the Lancones basin showing the main tectonic units and the spatial distribution of units sampled for isotopic analysis (including this study and data available from the literature; see references in the text). Chapter 5 Page 177 Figure 5.9. Thorogenic (A) and Uranogenic (B) Pb isotope diagrams for data from the Lancones Basin and fields for Pb isotope signatures of various tectonic units of the northern Andes or Perú and Ecuador. Data for ’Cretaceous platform sedimentary rocks’ and ‘continental crust (Olmos Complex)’ from the central Andes, Perú, after Macfarlane et al. (1990) and Macfarlane (1999). Data for ‘Jurassic‐Cretaceous metasedimentary rocks’ and ‘Jurassic‐Cretaceous MORB’, from Ecuador, after Bosch et al. (2002). Data for ‘Coastal Batholith’ from Mukasa (1986) and this study. Data for ‘East Pacific MORB’, East Pacific Rise, from Sun (1980). All data points are from this study. Symbols for rocks and sulphide samples and fields for Pb provinces are given in inset boxes in B. Chapter 5 Page 178 Figure 5.10. Rb‐Sr isotope plots for volcanic rocks of the Lancones Basin. A. Sr versus 87Sr/86Srί. B. 87Sr/86Sr versus 87Rb/86Sr. The line shown is defined by the felsic volcanic rocks only (n=6) and yields an errorchron of ~65 Ma. Chapter 5 Page 179 Figure 5.11 ‐ 143Nd/144Nd versus 87Sr/86Sr for volcanic rocks of the Lancones basin. A. Compared to fields for regional geologic units. B. Enlargement of data shown in A. Fields are given in the legend in A and symbols for both plots are shown in B. Data sources as per Fig. 5.9. Chapter 5 Page 180 Figure 5.12. Geochemical discrimination diagrams for volcanic rocks from the Lancones Basin. A. εNd versus Th/Yb. B. 206Pb/204Pb versus Th/Yb. Chapter 5 Page 181 Figure 5.13. Uranogenic Pb isotope diagram showing isotope compositions for mafic and felsic volcanic rocks at Tambogrande, as well as ore mineral isotope compositions. Volcanic rock data is conceptually time‐adjusted to 100 Ma based on Stacey and Kramers (1975) crustal growth curve with µ=9.85. Mixing lines are shown and labeled. Chapter 5 Page 182 Figure 5.14 ‐ Conceptual magmatic‐hydrothermal model relating petrogenesis of bimodal mafic‐felsic volcanic rocks at Tambogrande to hydrothermal system that formed VMS deposits (adopted from the petrogenetic model of Hart et al. (2004) for FII‐FIII felsic volcanic rocks). Depth of magma chamber is suggested by geochemical data which indicate partial melting or fractionation with amphibole ± pyroxene and plagioclase at shallow crustal depths. Isotope data support partial melting models as felsic volcanic rocks are substantially different isotopically as compared to basalts, and, yield more heterogeneous isotopic results. Hydrothermal leaching of metals, which is limited by the depth of fracture permeability, did not penetrate crustal rocks that were responsible for the unique isotope values (high 206Pb/204Pb) in the VMS‐associated felsic rocks. Chapter 5 Page 183 Table 5.1. Sample location data, approximate age and rock descriptions. All samples are from diamond drill core except for LW002. Coordinates are in map projection WGS 84, UTM Zone 17 Southern Hemisphere. Sample LW002 SU‐2 169.5 00TG1‐219 69.5 99‐TG3‐048 426.35 99TG3‐006 86.5 99TG3‐030 31.75 99TG3‐023 132.65 99TG3‐041 101 00TG1‐351 112 99TG1‐111 83.4 99TG3‐019 365.5 99‐TG3‐048 337.5 Easting Northing 571930 9475394 549200 9455650 573816 9454987 573612 9453492 573746 9453675 573815 9453628 573823 9453765 573823 9453915 573691 9455166 573696 9454532 573434 9453030 573612 9453492 Formation Cerro El Ereo Cerro El Ereo Cerro San Lorenzo Cerro San Lorenzo Cerro San Lorenzo Cerro San Lorenzo Cerro San Lorenzo Cerro San Lorenzo Cerro San Lorenzo Cerro San Lorenzo Cerro San Lorenzo Cerro San Lorenzo Lithology basalt basalt footwall basalt footwall basalt hanging wall basalt hanging wall basalt hanging wall dacite hanging wall dacite rhyolite dyke rhyolite dyke syn‐mineralization rhyolite syn‐mineralization rhyolite Age Ma 91‐100 91‐100 >104.8 >104.8 < 104.8 < 104.8 < 104.8 < 104.8 100.2 100.2 104.8 104.8 Comments plagioclase porphyritic flow plagioclase porphyritic flow autobreccia (clast sample); sericite‐chlorite alteration basalt flow; sericite‐chlorite alteration feldspar‐augite phyric flow feldspar‐augite phyric flow feldspar phyric flow feldspar phyric flow feldspar‐quartz porphyry dyke feldspar‐quartz porphyry dyke aphyric flow aphyric flow Chapter 5 Page 184 Table 5.2. Pb, Nd and Sr isotope data from volcanic rocks associated with VMS deposits at Tambogrande. ‘Initial’ isotope values are calculated to 100 Ma. Sample LW002 SU‐2 169.5 00TG1‐219 69.5 99‐TG3‐048 426.35 99TG3‐006 86.5 99TG3‐030 31.75 99TG3‐023 132.65 99TG3‐041 101 00TG1‐351 112 99TG1‐111 83.4 99TG3‐019 365.5 99‐TG3‐048 337.5 Pb ppm 5 <2 <2 3 <2 <2 <2 <2 <2 <2 <2 <2 U ppm <0.5 n/a n/a n/a n/a n/a n/a n/a n/a n/a n/a n/a Th ppm 0.33 0.22 0.25 0.65 1.04 0.89 4.12 4.66 8.96 8.27 4.17 3.21 206
Pb/204Pb 19.0147 18.8013 18.9723 19.0173 19.3685 18.9365 19.6574 20.3492 19.8028 19.6391 20.0273 20.1869 207
Pb/204Pb 15.6985 15.6518 15.6835 15.6861 15.6860 15.6608 15.7143 15.7463 15.7158 15.7056 15.7333 15.7380 208
Pb/204Pb 38.9072 38.6763 38.7822 38.8287 38.8556 38.6997 39.4066 39.7863 39.6656 39.4937 39.7836 39.4124 Sm Nd ppm ppm 1.17 3.38 2.38 7.23 1.22 3.26 2.08 6.04 2.54 8.68 2.61 9.52 4.30 16.08 4.32 16.62 5.28 21.45 4.01 16.85 6.72 22.93 5.76 18.72 Rb Sr 87
Nd/144Nd εNdί ppm ppm Sr/86Sr 0.512841 3.82 3 190 0.705667 0.512989 6.83 15 234 0.704872 0.512961 5.93 17 111 0.705865 0.512899 4.95 7 150 0.705400 0.512867 4.73 11 483 0.704370 0.512903 5.57 33 350 0.704063 0.512769 3.01 20 61 0.705722 0.512766 3.01 9 94 0.705254 0.512706 1.95 13 143 0.705437 0.512698 1.85 26 117 0.705501 0.512850 4.40 11 57 0.705621 0.512917 5.59 3 60 0.704778 143
87Sr/86Sr
n/a = not analyzed Chapter 5 ί
0.705602
0.704608
0.705235
0.705208
0.704276
0.703675
0.704374
0.704860
0.705063
0.704587
0.704828
0.704572
Page 185 Table 5.3. Pb isotope compositions of Tambogrande ore deposits and post‐mineralization intrusive phases in the Lancones basin. Sample B5 005‐542 TG1 028 20.4 TG1 237‐79.2 TG1 352‐94.5 TG3 009‐216 TG3‐037‐217 Easting Northing 569883 9442360 573638 9454725 573460 9455147 573455 9455227 573357 9453138 573822 9453844 LW‐31 484306 9426178 LW‐06 559404 9481200 LW‐36 LW‐85 567854 581496 9479640 9485511 rock type massive sulphide massive sulphide massive sulphide massive sulphide massive sulphide massive sulphide Muscovite‐bearing granodiorite, Amotape Terrane. Las Lomas batholith; hornblende granodiorite Las Lomas batholith; hornblende granodiorite porphyry Las Lomas batholith; hornblende granite Age Ma ~104 ~104 ~104 ~104 ~104 ~104 Mineral galena pyrite galena galena galena galena 206
207
208
18.8747 18.9211 18.9126 18.8928 18.8850 18.9325 15.6942 15.6951 15.6890 15.6726 15.6726 15.7025 38.7636 38.8118 38.7835 38.7227 38.7069 38.8548 ~230 plagioclase 18.5962 15.6682 38.5807 ~47 plagioclase 18.7234 15.6565 38.6497 ~52 ~80 plagioclase plagioclase 18.7142 18.9158 15.6577 15.6638 38.6888 38.9134 Pb/204Pb Pb/204Pb Pb/204Pb Chapter 5 Page 186 5.11 References Aspden, J. A., Bonilla,W., and Duque, P. 1995. The El Oro metamorphic complex, Ecuador: Geology and economic mineral deposits. Overseas Geology and Mineral Resources, no. 67, 63 p. Atherton, M.P., Pitcher, W.S., Warden, V. 1983. The Mesozoic marginal basin of central Perú. Nature, 350: 303‐306. Arculus, R.J., Lapierre, H., and Jaillard, E. 1999. Geochemical window into subduction and accretion processes; Raspas metamorphic complex, Ecuador. Geology, 27: 547‐550. Barrie, C.T. and Hannington, M.D. 1999. Classification of volcanic‐associated massive sulphide deposits based on host‐rock composition. 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Tosdal, R.M., Wooden, J.L., and Bouse, R.M. 1999. Pb isotopes, ore deposits, and metallogenic terranes. Reviews in Economic Geology, Volume 12, pp. 1‐28. Weis D., Kieffer B., Maerschalk C., Pretorius W. and Barling J. 2005. High‐precision Pb‐Sr‐Nd‐Hf isotopic characterization of USGS BHVO‐1 and BHVO‐2 reference materials. Geochemistry, Geophysics, Geosystems, DOI number 10.1029/2004GC000852. Winter, L.S., Tosdal, R., Franklin, J.M., Tegart, P. 2004. A Reconstructed Cretaceous Depositional Setting for Giant Volcanogenic Massive Sulfide Deposits at Tambogrande, Northwestern Perú. In Special Publication 11 ‐ Andean Metallogeny: New Discoveries, Concepts, and Updates, Edited by R.H. Sillitoe, J.Perello, and C.Vidal. Society of Economic Geologists, pp. 319‐340. Vergara, M., Levi, B., Nyström, J.O., Cancino, A. 1995. Jurassic and Early Cretaceous island arc volcanism, extension, and subsidence in the Coast Range of central Chile. Geological Society of America Bulletin, 107: 1427‐1440. Chapter 5 Page 191 Chapter 6. Summary, Discussion and Unresolved Questions 6.1 Summary and Main Conclusions The Cu‐Zn‐Au‐Ag volcanogenic massive sulphide (VMS) deposits at Tambogrande, are one of most significant metal accumulations of their type globally. The deposits formed in the Cretaceous Lancones basin of northwestern Perú, the result of an early Andean continental margin arc‐rift system. The elucidation of the timing of VMS formation, as well as the tectonic, regional geological, and paleo‐deposition setting are the salient goals of this study. A multidisciplinary approach was employed for this research project, which included mapping, three‐dimensional paleo‐seafloor reconstructions, geochronology, lithogeochemistry and isotope geochemistry. Using this broad scope, the ultimate goal of this research is to develop a comprehensive understanding of the genesis of the Tambogrande deposits, and to apply this knowledge to generic genetic models for VMS deposits. The results are summarized as follows: 1) The tectonic setting of the Lancones basin is a subduction‐related, continental margin arc‐rift system where arc volcanic rocks were deposited in segmented basins from Ecuador to southern Perú. The timing of volcanism within the Lancones basin ranges from a minimum of ~105 to 91 Ma (Chapter 2) and coincides broadly with the opening of the South Atlantic Ocean in the late stages of Gondwana break‐up (Scotese, 1991). The timing of volcanic arc formation is preceded by the accretion of the Amotape terrane to the Perú‐Ecuador continental margin in Early Cretaceous times (~132‐110 Ma; Arculus et al., 1999; Bosch et al., 2002). The accretion of this continental block precedes the initiation of the Lancones basin volcanism and therefore may have induced a new subduction system leading to the development of the volcanic arc. Slab rollback caused attenuation and rifting in the overriding plate led to the opening of the Lancones basin (Chapter 4). The final break‐up of Gondwana Chapter 6 Page 192 and westward drift of South America may have altered the dynamic along the western continental margin, limiting further expansion of the Lancones basin as the South American plate kept pace with the retreating arc (Soler and Bonhomme, 1990). The combination of a terrane accretion event and major plate tectonic reorganizations are the key factors leading to the development of the Lancones basin, and ultimately to the formation of VMS deposits at Tambogrande. Analogous modern ensialic arc rift settings include the Manus Basin (Taylor and Martinez, 1996) and Bransfield Strait (Lawver et al., 1995; González‐Casado et al., 2000). The Bransfield Strait may be the best analogue, and is an example of a marginal basin undergoing extension with mafic dominated, bi‐modal volcanism and has only minor MORB volcanism as part of incipient back‐arc spreading (Keller et al., 2002). 2) Volcanism in the Lancones basin is characterized by a diverse assemblage of marine volcanic and volcaniclastic rock facies that can be grouped into two major phases (Chapter 2). Phase 1, ca. 105 ‐ 100 Ma, comprises a mafic‐dominated volcanic stratigraphy represented by relatively deep water pillow lava flows and much less common felsic volcanic rocks. Felsic volcanic rocks represent a minor component during phase 1 and sedimentary rocks, other than minor pelagic sedimentary rocks, are rare. Phase 2, ca. 99 ‐ 91 Ma, is represented by relatively shallow water facies with a greater abundance of felsic pyroclastic, volcaniclastic and sedimentary rocks. Four formations are defined in this study, including the Cerro San Lorenzo Formation (Phase 1), and the Cerro El Ereo, La Bocana and Lancones formations (Phase 2). VMS deposits are associated with rhyolite flow dome complexes of the Cerro San Lorenzo Formation only and are confined to a narrow time‐stratigraphic horizon (Chapter 3). Chapter 6 Page 193 3) The paleo‐depositional setting of the Tambogrande VMS deposits (Chapter 3) is characterized by the development of a subaqueous, ~250 m‐thick rhyolite lava flow‐dome complex built upon a basaltic volcanic sequence. The deposits have low aspect ratios (average width/height) of approximately 2:5 and formed in deep, steep‐sided sub‐basins within the paleoseafloor, nested within the rhyolite complex. Massive sulphide lenses are up to 285 m thick and have an internal architecture characterized by a Cu‐rich base, pyrite‐rich core, and Zn‐
(Ba, Pb)‐rich upper margins, along with an underlying stockwork sulphide zone (Tegart et al., 2000). Both syn‐mineralization basalt and syn‐mineralization rhyolite flows were erupted episodically during VMS deposition, though these flows were not sufficient to completely bury or disrupt the hydrothermal circulation system. Northwest‐trending, syn‐volcanic structures at TG1 controlled basin development, volcanism and, ultimately, VMS deposit formation. The geometry of the sub‐basins was an important factor controlling the geometry and size of the deposits as they served as efficient ‘traps’ for depositing and preserving the massive sulphide deposits. In contrast to most other ancient ‘giant’ VMS deposits, models for the genesis of sulphide deposition at Tambogrande do not invoke substantial sub‐seafloor replacement of pre‐existing volcanic or sedimentary rocks. The focusing of mineralizing hydrothermal fluid through growing sulphide mounds within deep seafloor depressions resulted in efficient deposition and is considered key to forming and preserving the large accumulations of sulphide at Tambogrande. 4) A regional evaluation of the igneous petrogenesis of the bimodal, mafic‐dominated volcanic rocks of the Lancones Basin suggests magma genesis occurred in a supra‐subduction zone setting and was dominated by mantle sources, though continental basement was variably involved. All mafic volcanic rocks of the Lancones basin display prominent negative Nb Chapter 6 Page 194 anomalies on primitive mantle‐normalized trace element plots, similar to arc magmatic rocks (Chapter 4). The island arc geochemical affinities are attributed to the development of the extensional volcanic arc within relatively thin continental crust, unlike the modern Andes. This interpretation is consistent with geophysical modeling of the western Peruvian margin (Couch et al., 1981; Jones, 1981). No ‘non‐arc’ volcanic rocks (e.g., MORBs) were identified, indicating that back‐arc spreading did not occur within the Lancones basin. Therefore, the Lancones basin is interpreted to have formed in a rifted arc. However, Cretaceous intra‐continental alkaline volcanic rocks are preserved in the Eastern Cordillera from Ecuador to Argentina and represent the associated back‐arc (Soler and Bonhomme, 1990; Viramonte et al., 1999; Barragán et al., 2005). Back‐arc volcanism, however, did not evolve to sea‐floor spreading. Phase 1 and phase 2 volcanic events reflect different depositional environments (Chapter 2) and also yield subtly different geochemical signatures (Chapter 4) which can be used to interpret the tectonic evolution of the Lancones basin. Phase 1 Cerro San Lorenzo Formation basalt displays relatively high LREE/HREE ratios (i.e., La/Yb) and Zr/Y and have more calc‐alkaline affinities than the relatively tholeiitic Cerro El Ereo and La Bocana formations. These subtle lithogeochemical variations suggest progressive partial melting and depletion of the mantle wedge. In addition, phase 1 mafic volcanic rocks are relatively primitive (i.e., higher Mg) compared with phase 2 mafic rocks and were probably erupted through relatively thin crust. The more basaltic‐andesite dominant compositions of phase 2 formed in a maturing arc with thickened crust (Green, 1980; Gill, 1981). Phase 1 primitive lavas erupted during a period of extension, attenuated crust, basin subsidence, and a high geothermal gradient. Pb‐Nd‐Sr isotope systematics for mafic rocks indicate juvenile sources consistent with suprasubduction zone (mantle–wedge) derivation (Chapter 5). All felsic volcanic rocks analyzed Chapter 6 Page 195 in the study are subalkaline, M‐type, with negative Nb and Ti and positive Zr anomalies (Chapter 4). These felsic rocks have a basaltic parentage that was obtained either directly or indirectly, but exhibit variable radiogenic isotope signatures indicative of interaction continental basement rocks. This suggests these felsic rocks were generated within, or at least modified by, upper crustal material. Proterozoic zircon inheritance exhibited by some felsic volcanic rocks (Chapter 2) also supports the conclusion that continental crust influenced felsic magma genesis. The phase 1, VMS‐associated rhyolites at Tambogrande have low to moderate HFSE contents, flat REE patterns and tholeiitic‐transitional affinities characteristic of FII‐type rhyolites. Phase 2 rhyolites are relatively more depleted and are akin to FIV‐type felsic volcanic rocks. The phase 1 rhyolites originated from high‐temperature partial melts generated at shallow crustal levels. The petrogenetic assemblage of primitive mafic volcanic rocks and FII‐FIII rhyolites is an important feature of phase 1 and highlights the (i) high‐temperature magmatism and elevated geothermal gradient and (ii) extensional setting, both of which are critical to the initiation of a robust hydrothermal circulation system and VMS formation. 5) Lead isotope systematics (Chapter 5) provide a constraint on metal sources and therefore aid in the understanding of the hydrothermal system responsible for VMS formation at Tambogrande. The uniformity in Pb isotope compositions of massive sulphides from deposits TG1, TG3 and B5 suggests homogenization of Pb in the hydrothermal system. Massive sulphide Pb isotope compositions define a steep array on Pb isotope plots and are much less radiogenic than continental crust‐like felsic volcanic rocks, but very similar to basalt Pb isotope values. The spatially and temporally associated felsic volcanic rocks and VMS deposits at Tambogrande are genetically linked due to anomalous heat input in the upper crust, i.e., underplating of arc magmas. The emplacement of a mafic intrusion to sub‐crustal levels acted Chapter 6 Page 196 as a thermal driver for the hydrothermal system. In addition, this anomalous thermal input into the upper crust also caused partial melting of the juvenile mafic volcanic rocks. The resultant felsic partial melts were contaminated by continental basement rocks and erupted as felsic volcanic rocks. Although the felsic rocks may have played a role in heat transfer and enhanced the robustness of the hydrothermal system, the role of these rocks is considered largely passive in the formation of the Tambogrande VMS deposits. Metals were leached almost exclusively from mafic footwall strata. In summary, the key contributions of this research project are summarized as follows: (1) the thesis provides comprehensive baseline documentation of an economically significant VMS district that has not hitherto been studied in any significant detail. The work here is built on a foundation of first understanding the regional geology, deposit‐scale volcanic setting, and deposit architecture; (2) research on the Lancones basin has helped fill a regional void in the understanding of the geology of Cretaceous volcanism and tectonics in the northwest Perú Andean segment. The data also provide a better understanding of VMS metallogenesis within marginal basins of Perú; (3) the reconstruction of the paleo‐depositional environment for massive sulphide mineralization at Tambogrande provides a well constrained model for the seafloor depositional setting of a giant VMS deposit. As many ancient VMS deposits are often too deformed to construct any meaningful models, this study provides a unique perspective; (4) the thesis provides a synopsis of a continental‐margin VMS setting with strong constraints on the relationship to global plate tectonics; and (5) the research helps constrain models for VMS genesis in terms of the relationship to associated felsic volcanic rocks, metal sources and hydrothermal system generation. Chapter 6 Page 197 6.2 Discussion and Ideas 6.2.1 Timing and Tectonic Setting A peri‐cratonic extensional setting for the Lancones Basin is implicit because the present day volcanic sequence is ~150 km wide, with continental crust preserved adjacent to both margins. Further, it is also underlain in part by continental basement as evidenced by radiogenic isotopes and zircon inheritance. The role of extensional geodynamics in the formation of VMS deposits has been frequently emphasized in ancient VMS settings (e.g., Sillitoe, 1982; Cathles, 1983; Swinden, 1991; Lentz, 1998; Hart et al., 2004) and observed in modern VMS environments (Hannington et al., 2005). Although the kinematics of such modern environments are relatively well understood, though often debated (e.g., Bransfield Strait; Lawver et al., 1995; Barker and Austin, 1998; González‐Casado et al., 2000; Fretzdorff et al., 2004), the forces that controlled extension in ancient marginal arcs remain less well understood. Extension can be accommodated in various ways. Extension at Tambogrande was probably associated with slab rollback and oceanwards retreat of the overlying arc due to subduction of relatively old and dense oceanic material during Albian times (Soler and Bonhomme, 1990). Alternatively, and considered less likely due to the volume of volcanic material erupted, transpressional or transtensional tectonics may also have controlled the development of the Lancones Basin as a pull‐apart basin. The timing of VMS formation at Tambogrande is temporally linked with major plate tectonic reorganizations, specifically, the final breakup of Gondwana and the Albian rifting of the equatorial South Atlantic (Scotese, 1991). The extensional basin tectonics and VMS‐setting was terminated with the ultimate separation of South America from Africa and the onset of spreading in the South Atlantic, causing the rotation and westward drift of the South American continent and contractional tectonics along the western margin, although the exact kinematics Chapter 6 Page 198 are uncertain (Soler and Bonhomme, 1990). In general, there appears to be an association of VMS deposits with major plate re‐organizations elsewhere in the rock record (Barley and Groves, 1992; Groves and Bierlein, 2007). This conclusion is often made based on the shift from arc to non‐arc (i.e., MORB) volcanism, indicating the transition from arc rifting to active spreading (e.g., Swinden, 1991; Swinden et al., 1997; Wyman et al., 1999; Piercey, 2007). At Tambogrande, the transition was not from arc to non‐arc volcanism, rather a continuous arc volcanic sequence evolved from extensional to a more neutral and eventually contractional tectonic regime. VMS deposits are commonly associated with ‘aborted arc rifts’ (e.g., Kuroko; Cathles et al., 1983) which typify microplate tectonic regimes of the SW Pacific Manus Basin (Martinez and Taylor, 1996). 6.2.2 Where’s the Intrusion? VMS genetic models advocate an igneous heat source to drive a hydrothermal circulation system and possibly to contribute metals (Lydon, 1988; Large, 1992; Galley, 1993; Franklin, 1995; Ohmoto, 1996; Hart et al., 2004; Franklin et al., 2005) . This source of heat is generally considered to be new magma input to the upper crust. Heat‐ and fluid‐flow modeling indicate that substantial thermal input into the upper crust is required in order to sustain a hydrothermal convection system capable of producing giant massive sulphide deposits (Cathles et al., 1981; Cathles et al., 1999; Barrie et al., 1999). Furthermore, the thermal input required is most likely accommodated by mantle‐derived intrusions (mafic, or ultramafic) emplaced within the upper crust (Barrie et al., 1999). That bimodal‐volcanism did not bury the deposits at Tambogrande but was episodic during deposition of the massive sulphides suggests stable and likely periodic replenishment of the magma chamber(s). The presence of petrochemically variable, volumetrically minor felsic volcanic rocks intercalated with massive sulphide suggests a complex magmatic system erupted small batches of magma during and after massive sulphide Chapter 6 Page 199 deposit formation, likely the result of a long‐lived magma chamber and area of elevated geothermal gradient. The volcanic stratigraphy at Tambogrande is near flat lying and not deeply eroded, therefore, such intrusions are not likely exposed in the region. 6.3 Outstanding Issues and Directions for Future Research 6.3.1 At Tambogrande and in the Lancones Basin • Was Tambogrande related to a long‐lived hydrothemal system? An attempt to date the immediate footwall and hanging wall rocks at Tambogrande was not successful due to the lack of zircons in most of the felsic volcanic rocks, likely due to the inferred high temperatures at eruption (Chapter 2). For comparison, Bleeker et al. (1999) suggest the Kidd Creek ore body formed during a prolonged, but episodic, period of hydrothermal activity spanning ~ 5 Ma. How else can these (and other) stratiform volcanic‐hosted deposits be dated? •
What is the maximum age of volcanism in the Lancones Basin? When did the basin open? Do granitic plutonism and phase 2 volcanism temporally overlap in the Lancones basin? In the Western Peruvian Trough? •
Where is the northern terminus of the arc in Ecuador? In Columbia? •
Based on known resources, the Tambogrande district contain approximately 3.5 million tonnes of Cu metal in the three massive sulphide deposits. In a hypothetical metal mass balance model the footwall basalt rocks, which served as the source of the metals, contain a nominal 100 ppm Cu and have a density of 2.7 g/cm3, and Cu is leached to the hydrothermal fluid at 10% efficiency rate (cf. Barrie et al., 1999). In modern settings, metal loss from hydrothermal vents to ocean water is 95% to >99% (Converse et al., 1984; Feely et al., 1994). This explains why most mound style VMS deposits forming on the ocean floor are less than 1 Mt, whereas modern subseafloor replacement systems Chapter 6 Page 200 tend to be substantially larger (Goodfellow and Zierenberg, 1999). If depositional efficiency at the vent site during the formation of the Tambogrande deposits was 5% like modern seafloor mound‐type settings, the volume of footwall rocks affected by the hydrothermal system would be ~1300 km3. Where in the rock record or in modern systems do such giant hydrothermal systems exist? Chapter 3 suggests sulphide deposition at Tambogrande was different from many mound type deposits and was quite efficient, though this is difficult to quantify precisely. Was deposition exceptionally efficient? If such large‐scale hydrothermal systems cannot be demonstrated in the rock record or in modern settings, the probability is enhanced of exceptional depositional efficiency from an ‘average’ hydrothermal system at Tambogrande. Many of the outstanding questions at Tambogrande will only be answered by additional drilling, geophysical surveys, and generally with more exploration producing more data, such as: •
What are the key structural controls that link the deposits? Are TG1, TG3 and B5 sites of NNW directed pull‐apart sub‐basins linked to a NNE striking transtensional fault? •
What is the ultimate sulphide tonnage of the district? VMS districts usually contain a range of deposit sizes with a range of metal grades. The three deposits discovered at Tambogrande so far are very uniform in their grade and tonnage. Are there any smaller or ‘normal’ size deposits? Are there higher grade deposits? 6.3.2 In the Western Peruvian Trough Massive sulphide deposits throughout the Cretaceous Western Peruvian Trough of Perú are hosted within the Casma Group (Vidal, 1987). However, the age of some VMS deposits (i.e., Chapter 6 Page 201 Perúbar) of the Western Peruvian Trough are reported as latest Late Cretaceous (ca. 69 Ma) and suggested to have formed in pull‐apart basins related to wrench tectonics (Polliand et al., 2005). The geochronology of the Lancones basin presented herein shows the Tambogrande VMS deposits formed ca. 105 Ma. In comparison, based on fossil ages, Casma Group volcanism is limited to the middle Albian (Myers, 1974; Cobbing et al., 1981). Therefore, the Cerro San Lorenzo formation is the temporal and stratigraphic equivalent to the VMS‐hosting volcanic successions of the Casma Group. Excluding Perúbar, VMS deposits of the Western Peruvian Trough and Lancones basin represent a major VMS‐forming event in Albian times along marginal rift basins of the central Andes. The following are avenues for additional research: •
Limited geochronological work has been undertaken in the Western Peruvian Trough. What is the duration of arc volcanism in each of these basins? Oblique collisional zones can result in hinged rift basins. Are there variations that might indicate the direction of rift propagation, or rift termination? Did VMS deposits in the Western Peruvian Trough and Lancones basin form at the same time? Are there multiple episodes of VMS formation? How does these relate with respect to the precise timing of formation plate tectonic reorganization in the South Atlantic? •
The Lancones Basin and Western Peruvian Trough are pericratonic arc‐related basins, but the influence of continental crust is variable as evidenced in the isotopic data for mafic and felsic volcanic rocks in the Lancones basin. What are the variations in continental crust involvement with respect to magmatism along the length of the arc? Are there both oceanic‐dominated and continental arc‐dominated domains? Additional geochemical and isotopic data are required in these other basins to address this issue. Chapter 6 Page 202 •
Extension and rifting of an arc complex are indicated for the Lancones basin, but no subsequent seafloor spreading is known to have occurred (i.e., absence of mid‐ocean ridge basalt, or MORB). The Western Peruvian Trough may have MORB‐type basalt in upper sections (Atherton et al., 1983), indicating that back‐arc spreading evolved to a greater degree than in the Lancones Basin. Otherwise, back‐arc rifting is represented by alkaline volcanism farther east. Additional lithogeochemical and isotopic studies are required to document the petrochemistry of these sequences. •
Why are there no giants in Western Peruvian Trough? Cerro Lindo is by far the largest at 42 Mt (Steinmüller, 2000). Does this suggest something unique about the massive sulphide forming environment in the Lancones basin segment of the Peruvian marginal basins? 6.3.3 Globally • Most VMS tectonic setting classifications are based on volcanic rock compositions, facies analysis and geochemistry. The Tambogrande example illustrates that well constrained plate tectonic models and paleogeographic reconstructions are necessary to fully appreciate the tectonic setting of many VMS districts. Are continental margin arc volcanic settings underestimated in terms of importance in the rock record? In deformed terranes are they too often incorrectly interpreted as allocthonous oceanic arc terranes? •
VMS deposits are often confined to specific time stratigraphic horizons. The timing relative to various magmatic and tectonic processes is an area where a better understanding is required. Also, the longevity of these systems remains poorly understood. Chapter 6 Page 203 •
The genesis of giant VMS systems remains enigmatic, with the world’s largest deposits yielding variable compositions and occurring in various host rocks in different settings and within rocks of all ages. Different controls on the genesis of giant deposits have been demonstrated though much work is required before these criteria can have practical applications to mineral exploration. Chapter 6 Page 204 6.4 References Arculus, R.J., Lapierre, H., and Jaillard, E. 1999. Geochemical window into subduction and accretion processes; Raspas metamorphic complex, Ecuador, Geology, 27: 547‐550. Atherton, M.P., Pitcher, W.S., Warden, V. 1983. The Mesozoic marginal basin of central Perú. Nature, 350: 303‐306. Barker, D. H. N., and Austin, Jr. J. A. 1998. Rift propagation, detachment faulting, and associated magmatism in Bransfield Strait, Antarctic Peninsula, Journal of Geophysical Research, 103(B10): 24,017‐24,043. Barley, M. E., and Groves, D. I. 1992. Supercontinent cycles and the distribution of metal deposits through time. Geology, 20: 291‐294. Barragán, R., Babya, P., and Duncan, R. 2005. Cretaceous alkaline intra‐plate magmatism in the Ecuadorian Oriente Basin: Geochemical, geochronological and tectonic evidence. Earth and Planetary Science Letters, 236: 670‐690. Barrie, C.T., Cathles, L.M., and Erendi, A. 1999. Finite element and fluid‐flow computer simulations of a deep ultramafic sill model for the giant Kidd Creek volcanic‐associated massive sulphide deposit, Abitibi subprovince. Economic Geology Monographs, 10: 529‐
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324. Green, T.H. 1980. Island arc and continent‐building magmatism; a review of petrogenic models based on experimental petrology and geochemistry. Tectonophysics, 63: 367‐385. Groves, D. I. and Bierlein, F. P. 2007. Geodynamic settings of mineral deposit systems. Journal of the Geological Society, 164: 19‐30. Chapter 6 Page 206 Hannington, M.D., de Ronde, C.E.J., and Petersen, S. 2005. Sea‐floor tectonics and submarine hydrothermal systems. In 100th Anniversary Volume, 1905‐2005. Edited by Jeffrey W. Hedenquist, J.F.H. Thompson, R.J. Goldfarb, and J.P. Richards. Society of Economic Geologists, pp. 111‐142. Hart, T.R., Gibson, H.L., and Lesher, C.M. 2004. Trace element geochemistry and petrogenesis of felsic volcanic rocks associated with volcanogenic massive Cu‐Zn‐Pb sulfide deposits. Economic Geology, 99: 1003‐1013. Jones, P.R. 1981. Crustal structures of the Perú continental margin and adjacent Nazca Plate, 9 degrees S latitude. Geological Society of America, Memoir 154, pp. 423–443. Keller, R.A., Fisk, M.R., Smellie, J.L., Strelin, J.A., Lawver, L.A. 2002. Geochemistry of back arc basin volcanism in Bransfield Strait, Antarctica; subducted contributions and along‐axis variations. Journal of Geophysical Research, 107: 10.1029/2001, JB000444, 17p. Large, R.R. 1992. Australian volcanic‐hosted massive sulfide deposits: Features, styles, and genetic models. Economic Geology, 87: 471–510. Lawver, L. A., Keller, R. A., Fisk, M. R. and Strelin, J. 1995. Bransfield Strait, Antarctic Peninsula: Active Extension Behind a Dead Arc. In Back‐Arc Basins: Tectonics and Magmatism. Edited by B. Taylor. Plenum Publishing Corp., New York, pp. 315–342. Lentz, D.R. 1998. Petrogenetic evolution of felsic volcanic sequences associated with Phanerozoic volcanic‐hosted massive sulphide systems; the role of extensional geodynamics. Ore Geology Reviews, 12: 289‐327. Lydon, J.W. 1988. Volcanogenic massive sulphide deposits, part 2, genetic models; Geoscience Canada, 15: 43‐65. Martinez, F., and Taylor, B. 1996. Fast backarc spreading, rifting, and microplate rotation, between transform faults in the Manus Basin, Bismarck Sea. Marine Geophysics Research, 18: 203‐224. Myers, J.S. 1974. Cretaceous stratigraphy and structure, Western Andes of Perú between latitudes 10o‐10o30 . The American Association of Petroleum Geologists Bulletin, 58: 474‐
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Magmatic and tectonic evolution of the Cambrian‐Ordovician Laurentian margin of Iapetus; geochemical and isotopic constraints from the Notre Dame subzone, Newfoundland. In the nature of magmatism in the Appalachian orogen. Edited by K. Sinha, J.B. Whalen, and J.P. Hogan. Geological Society of America, Memoir 191, pp. 337‐365. Taylor, B., and Martinez, F. 2003. Back‐arc basin basalt systematics. Earth and Planetary Science Letters, 210: 481‐497. Tegart, P., Allen, G., Carstensen, A. 2000. Regional setting, stratigraphy, alteration and mineralization of the Tambo Grande VMS district, Piura Department, Northern Perú. In VMS Deposits of Latin America. Edited by R. Sherlock and M.A.V. Logan. Geological Association of Canada, Mineral Deposits Division, Special Paper No.2. pp. 375‐405. Vidal, C.E. 1987. Kuroko‐type deposits in the Middle‐Cretaceous marginal basin of central Peru. Economic Geology, 82: 1409‐1430. Viramonte, J. G., Kay, S. M., Becchio, R., Escayola, M., and Novitski, I. 1999. Cretaceous rift related magmatism in central‐western South America. Journal of South American Earth Sciences, 12: 109‐121. Chapter 6 Page 208 Wyman, D.A., Bleeker, W., and Kerrich, R. 1999. A 2.7 Ga plume, proto‐arc, to arc transition and the geodynamic setting of the Kidd Creek deposit: evidence from precise ICP MS trace element data. In The Giant Kidd Creek volcanogenic massive sulfide deposit, western Abitibi subprovince. Edited by M.D. Hannington and T.C. Barrie. Economic Geology, Monograph 10, pp. 525–542. Chapter 6 Page 209 Appendix A. U‐Pb Zircon Sample Preparation, Analysis and Additional Data A.1 Methodology Zircon was separated from approximately 10‐15 kg of fresh rock material was used of each sample. Samples for U‐Pb dating were processed using a Rhino™ jaw crusher, a Bico™ disk grinder equipped with ceramic grinding plates, and a Wilfley wet shaking table equipped with a machined Plexiglass top, followed by conventional heavy liquids and magnetic separation using a Frantz™ magnetic separator. Samples were hand‐picked using a binocular microscope. Thermal Ionization Mass Spectrometry (TIMS) U‐Pb analyses were done at the Pacific Centre for Isotopic and Geochemical Research at the University of British Columbia. The methodology for zircon grain selection, abrasion, dissolution, geochemical preparation and mass spectrometry are described by Mortensen et al. (1995). All zircon fractions were air abraded (Krogh, 1982) prior to dissolution to minimize the effects of post‐crystallization Pb‐loss. Procedural blanks for Pb and U were 2 and 1 pg, respectively. Analytical data are listed in Table A1. Errors attached to individual analyses were calculated using the numerical error propagation method of Roddick (1987). Decay constants used are those recommended by Steiger and Jäger (1975). Compositions for initial common Pb were taken from the model of Stacey and Kramer (1975). All errors are given at the 2 sigma level. Sensitive High Resolution Ion Microprobe ‐ Reverse Geometry (SHRIMP) analysis was carried out Stanford University, Stanford, California. All samples were embedded in 25.4 mm epoxy rounds. Morphology and internal structure was examined by SEM using both BSE imaging and cathodoluminescence to view polished grain mounts (Fig. A1). Data is presented in table A2. Appendix A Page 210 Table A1. U‐Pb zircon analytical data obtained using ID‐TIMS method. 2
206Pb/238U4 (± % 1σ ) Sample 1
Description Wt (mg) U (ppm) Pb (ppm) 206Pb/204Pb 3
(meas.) total common Pb (pg) A: N2,+74 B: N2,+74 C: N2,+74 D: N2,+74 E: N2,+74 0.021 0.023 0.024 0.023 0.039 334 270 271 98 231 5.5 4.3 4.5 1.6 3.7 1653 909 385 624 1006 4 7 17 4 9 A: N1,74‐104 C: N1,‐74 D: N1,‐74 E: N1,‐74 0.032 0.027 0.028 0.032 280 404 295 294 4.4 6.5 4.7 4.6 843 570 698 749 10 18 11 12 A: N2,+74 B: N2,‐74 C: N2,‐74 D: N2,‐74 E: N2,‐74 F: N2,‐74 0.043 0.018 0.048 0.053 0.068 0.055 294 371 334 361 443 401 4.4 6.0 5.2 5.6 6.8 6.1 459 336 1139 405 642 915 25 19 13 44 42 22 A: N2,+74 B: N2,+74 C: N2,+74 D: N2,+104 E: N2,+74 F: N2,+74 Sample LW‐151 A: N2,+74 B: N2,+104 C: N2,<74 0.044 0.043 0.058 0.048 0.015 0.017 531 441 432 306 631 392 8.6 7.3 7.0 4.9 10.3 6.3 1189 1750 1274 1150 825 318 19 10 19 12 11 20 Sample 99TG‐1 0.01567(0.23) 0.01553(0.42) 0.01587(0.22) 0.01566(0.38) 0.01569(0.16) Sample LW‐010 13.1 0.01514(0.13) 14.7 0.01518(0.15) 14.2 0.01516(0.13) 13.8 0.01510(0.12) Sample LW‐077 14.2 0.01422(0.22) 16.7 0.01483(0.26) 15.1 0.01457(0.19) 14.8 0.01470(0.18) 15.1 0.01442(0.13) 14.6 0.01452(0.12) Sample LW‐086 14.9 0.01534(0.13) 15.1 0.01553(0.13) 13.9 0.01551(0.14) 11.3 0.01562(0.13) 15.1 0.01540(0.28) 15.4 0.01502(0.24) 0.018 0.052 0.060 489 486 303 10.8 5.8 19.0 2583 1417 1005 4 13 67 14.2 13.7 11.0 % 208Pb2 13.9 12.5 14.0 13.9 12.9 0.02069(0.19) 0.01147(0.18) 0.05982(0.14) 207Pb/235U4 (± % 1σ ) 207Pb/206Pb4 (± % 1σ ) 206Pb/238U age (Ma; ± % 2 σ) 207Pb/206Pb age (Ma; ± % 2σ ) 0.1043(0.50) 0.1028(0.57) 0.1090(0.73) 0.1085(1.24) 0.1040(1.00) 0.04826(0.42) 0.04799(0.66) 0.049840.62) 0.05029(1.17) 0.04805(0.96) 100.2(0.5) 99.3(0.8) 101.5(0.4) 100.2(0.8) 100.4(0.3) 112.0(20.0) 98.8(31.1) 187.4(28.9) 208.4(54.3) 101.8(45.6) 0.1001(0.68) 0.1004(1.50) 0.1002(1.03) 0.0998(0.57) 0.04795(0.62) 0.04797(1.43) 0.04796(0.98) 0.04794(0.51) 96.9(0.3) 97.1(0.3) 97.0(0.3) 96.6(0.2) 96.8(29.5) 97.5(67.9) 97.3(46.2) 96.5(24.2) 0.0938(2.30) 0.0974(0.87) 0.0961(0.37) 0.0972(1.29) 0.0952(1.33) 0.0958(0.78) 0.04783(2.19) 0.04764(0.75) 0.04781(0.29) 0.04795(1.21) 0.04788(1.27) 0.04786(0.74) 91.1(0.4) 94.9(0.5) 93.3(0.4) 94.1(0.3) 92.3(0.2) 92.9(0.2) 91.0(105.5) 81.5(35.6) 89.8(13.4) 96.6(57.2) 93.3(60.6) 92.4(35.0) 0.1034(0.86) 0.1027(0.60) 0.1027(0.81) 0.1053(0.31) 0.1013(0.61) 0.0993(2.91) 0.04889(0.82) 0.04794(0.56) 0.04802(0.77) 0.04888(0.24) 0.04771(0.50) 0.04795(2.79) 98.1(0.3) 99.3(0.3) 99.2(0.3) 99.9(0.3) 98.5(0.5) 96.1(0.5) 142.7(38.3) 96.5(26.4) 100.1(36.2) 141.9(11.4) 84.8(23.8) 96.7(132.4) 0.1808(0.29) 0.0785(0.39) 0.7176(0.28) 0.06337(0.21) 0.04963(0.33) 0.08700(0.20) 132.0(0.5) 73.5(0.3) 374.5(1.0) 720.8(8.8) 177.5(15.2) 1360.5(7.6) 1
N1,N2 = non‐magnetic at n degrees side slope on Frantz magnetic separator; grain size given in microns. radiogenic Pb; corrected for blank, initial common Pb, and spike 3
corrected for spike and fractionation 2
4
corrected for blank Pb and U, and common Pb Appendix A Page 211 Table A2. U‐Pb zircon analytical data obtained using the SHRIMP‐RG method. Meas 204
Pb/ 206
Pb Spot Name % err Meas 207
Pb/ 206
Pb % err Meas 208
Pb
/ 206
Pb % err Th ppm U ppm 232
/
Th
U Corr 206
Pb
238
/ U % err Rad 206
Pb (ppm) 238
207 corr 206
Pb /238U Age 1σ err 204 corr 207
Pb 206
/ Pb % err 204 corr 207
Pb
235
/ U 204 corr 206Pb
/238U % err % err err corr TG1‐136 ‐ dacite pebble breccias 1.1* 0.001146 65 .053 6.2 .082 8.9 39 129 0.26 .016 2.1 2.1 100.9 2.1 .0358 33.5 0.08 33.6 .0156 2.4 .073 2.1* 0.000000 0 .049 3.8 .168 3.0 158 287 0.50 .016 1.3 4.4 100.1 1.4 .0489 3.8 0.11 4.1 .0157 1.3 .330 3.1* 0.000453 47 .049 4.3 .152 3.6 117 227 0.46 .016 1.5 3.6 99.8 1.5 .0424 9.3 0.09 9.4 .0155 1.5 .163 4.1 0.000000 0 .050 4.6 .141 3.9 102 208 0.43 .017 1.6 3.5 106.4 1.7 .0500 4.6 0.11 4.8 .0167 1.6 .321 5.1 0.000071 52 .050 1.7 .246 1.0 1521 1807 0.76 .016 0.7 29.3 104.8 0.8 .0494 2.1 0.11 2.2 .0164 0.7 .334 6.1 0.000000 0 .051 3.2 .263 1.8 690 771 0.79 .016 1.4 12.5 103.1 1.5 .0506 3.2 0.11 3.5 .0162 1.4 .404 7.1 0.000574 55 .049 5.6 .200 2.6 218 329 0.58 .017 1.1 5.6 106.5 1.2 .0408 13.8 0.09 13.8 .0165 1.2 .090 8.1* 0.000000 0 .058 4.7 .140 3.8 110 228 0.43 .018 1.3 4.0 111.6 1.5 .0577 4.7 0.14 4.9 .0176 1.3 .271 9.1 0.000120 49 .050 2.7 .228 1.6 611 769 0.70 .016 0.9 12.6 104.1 0.9 .0487 3.3 0.11 3.5 .0163 0.9 .256 10.1 0.000291 49 .051 4.2 .148 3.1 153 309 0.44 .017 1.7 5.3 108.4 1.8 .0469 6.6 0.11 6.8 .0169 1.7 .247 11.1 0.000000 0 .045 5.3 .114 4.9 88 232 0.34 .016 1.4 3.7 102.0 1.5 .0452 5.3 0.10 5.5 .0159 1.4 .261 12.1 0.000703 56 .049 5.0 .141 7.7 119 253 0.42 .016 1.3 4.2 104.6 1.4 .0386 17.1 0.09 17.2 .0162 1.5 .086 LW‐016 ‐ rhyolite quartz porphyry dyke 1 64 .048 4.6 .215 2.9 127 197 0.66 .016 1.2 2.8 105.1 1.3 .0522 7.1 0.12 7.2 .0165 1.3 .175 2* 0.000329 38 .048 2.9 .283 1.7 417 495 0.87 .017 0.7 7.1 106.9 0.8 .0436 5.4 0.10 5.5 .0166 0.8 .142 3 0.000922 27 .067 3.8 .225 2.8 133 237 0.58 .016 1.1 3.3 101.1 1.2 .0536 8.5 0.12 8.6 .0159 1.2 .141 4 0.000127 40 .050 3.1 .338 1.7 483 483 1.03 .016 0.8 6.8 104.2 0.9 .0482 3.6 0.11 3.7 .0163 0.8 .221 63 .067 6.9 .231 4.9 43 86 0.52 .017 2.0 1.2 103.8 2.1 .0782 10.5 0.18 10.7 .0168 2.1 .200 6 0.000377 39 .052 4.7 .181 3.4 113 213 0.55 .016 1.2 3.0 103.0 1.3 .0466 7.1 0.10 7.2 .0161 1.3 .176 7 0.000523 40 .049 4.5 .190 3.2 143 251 0.59 .016 1.2 3.5 104.4 1.2 .0416 9.3 0.09 9.4 .0162 1.2 .130 3 .049 5.8 .249 3.5 139 179 0.80 .016 1.5 2.5 102.9 1.5 .0582 4.9 0.13 5.1 .0163 1.5 .284 0.000407 30 .048 3.2 .356 2.0 512 480 1.10 .016 1.0 6.7 103.4 1.0 .0418 5.8 0.09 5.9 .0160 1.0 .167 5 8 9 10* 0 .050 2.3 .414 1.1 1219 986 1.28 .017 0.6 14.2 107.0 0.6 .0496 2.3 0.11 2.3 .0168 0.6 .250 11 0.000000 61 .049 12.8 .147 8.5 24 57 0.43 .016 2.7 0.8 100.6 2.8 .0498 12.5 0.11 12.8 .0158 2.7 .212 12 2 .049 4.7 .265 2.8 253 326 0.80 .016 1.2 4.4 101.0 1.2 .0553 4.1 0.12 4.3 .0159 1.2 .274 Appendix A Page 212 Table A2 continued. Meas 204Pb/ 206Pb Meas 207
Pb/ 206
Pb % err % err Meas 208
Pb
/ 206
Pb % err Th ppm U ppm 232
Th /238U Corr 206
Pb /238U % err Rad 206
Pb (ppm) 207 corr 206
Pb /238U Age 1σ err 204 corr 207
Pb /206Pb % err 204 corr 207
Pb /235U 204 corr 206Pb /238U % err % err err corr LW‐013 ‐ rhyolite quartz porphyry dyke 1 0.000982 16 .048 4.3 .089 5.3 52 198 0.27 .015 1.5 2.6 96.6 1.4 .0334 11.0 0.07 11.1 .0148 1.5 .135 2a 0.000758 34 .047 3.0 .174 2.9 185 358 0.53 .016 1.1 4.9 102.4 1.2 .0352 12.4 0.08 12.4 .0158 1.2 .098 2b 0.000661 38 .050 3.0 .155 3.2 150 343 0.45 .015 1.1 4.5 97.8 1.1 .0404 10.5 0.08 10.5 .0151 1.2 .116 3* 0.001972 16 .051 4.5 .151 4.0 107 268 0.41 .017 1.4 3.9 106.5 1.6 .0209 32.2 0.05 32.3 .0161 1.6 .048 4 0.000201 67 .046 4.0 .091 5.2 57 198 0.30 .015 1.5 2.6 98.4 1.5 .0430 6.4 0.09 6.6 .0153 1.5 .231 5 6 0.000457 7 8* 0.012278 69 .046 2.8 .191 3.8 235 410 0.59 .016 1.2 5.5 100.7 1.2 .0396 12.6 0.09 12.7 .0156 1.3 .104 108 .049 1.7 .171 2.1 590 1160 0.53 .016 0.6 15.5 99.5 0.6 .0493 1.9 0.11 2.0 .0156 0.6 .318 9 .235 3.9 .601 4.5 120 273 0.46 .020 1.1 4.7 98.5 4.0 .0566 65.8 0.12 65.8 .0156 2.8 .042 42 .050 6.2 .135 4.5 74 198 0.39 .017 1.7 2.9 110.5 1.9 .0657 11.5 0.16 11.7 .0177 1.9 .164 9 0.001503 30 .048 3.7 .164 3.8 117 246 0.49 .016 1.4 3.5 104.6 1.5 .0253 31.0 0.06 31.0 .0159 1.6 .053 10 0.001397 28 .050 3.1 .201 2.9 168 307 0.56 .016 1.2 4.1 99.3 1.2 .0289 23.3 0.06 23.3 .0152 1.4 .059 11* 0.003831 24 .066 3.5 .156 4.7 124 274 0.47 .017 1.5 4.1 108.5 1.7 .0162 2.3 LW‐078 ‐ rhyolite boulder volcaniclastic 1* 0.000814 32 .050 4.6 .140 3.7 89 216 0.43 .015 1.2 2.8 95.4 1.2 .0374 12.5 0.08 12.6 .0147 1.3 .103 2 0.000802 27 .046 6.3 .090 6.0 34 125 0.28 .016 1.6 1.7 100.4 1.6 .0342 13.2 0.07 13.3 .0154 1.6 .122 56 .052 5.9 .128 5.1 47 150 0.32 .016 1.5 2.0 99.9 1.6 .0666 13.2 0.15 13.3 .0160 1.8 .138 0.000860 26 .058 4.9 .139 4.3 55 170 0.33 .016 1.4 2.3 98.4 1.4 .0450 10.1 0.10 10.2 .0153 1.4 .141 2 .051 6.4 .095 6.1 36 125 0.29 .016 1.6 1.7 99.4 1.7 .0543 5.9 0.12 6.2 .0157 1.6 .265 6 0.000416 67 .049 4.7 .128 3.9 77 214 0.37 .015 1.2 2.8 96.5 1.2 .0429 11.1 0.09 11.2 .0150 1.3 .118 7 0.000180 65 .048 3.6 .155 2.7 149 336 0.46 .016 0.9 4.5 100.5 0.9 .0450 5.5 0.10 5.5 .0157 0.9 .171 8 0.000032 1278 .059 6.3 .089 6.8 23 109 0.22 .015 1.7 1.4 97.6 1.7 .0585 11.9 0.12 12.1 .0155 1.9 .155 9 0.000418 47 .049 3.9 .133 3.2 114 287 0.41 .016 1.0 3.8 99.7 1.0 .0428 8.3 0.09 8.4 .0155 1.1 .129 3 4 5 10 11 12 0.000990 0.001679 44 .050 5.2 .144 4.1 67 161 0.43 .016 1.4 2.1 99.0 1.4 .0348 20.8 0.07 20.8 .0152 1.6 .076 46 .049 2.6 .177 1.9 344 647 0.55 .015 0.7 8.6 98.4 0.7 .0511 3.0 0.11 3.1 .0154 0.7 .222 23 .054 7.5 .104 9.4 21 73 0.29 .015 2.0 1.0 96.9 2.0 .0291 25.7 0.06 25.8 .0148 2.1 .082 Appendix A Page 213 Table A2 continued. Meas 204
Pb/ 206
Pb Spot Name Meas 207
Pb/ 206
Pb % err % err Meas 208
Pb
/ 206
Pb % err Th ppm U ppm 232
Th /238U Corr 206
Pb /238U % err Rad 206
Pb (ppm) 207 corr 206
Pb /238U Age 204 corr 207
Pb /206Pb 1σ err 204 corr 207
Pb /235U % err % err 204 corr 206Pb /238U % err err corr LW‐043 ‐ rhyolite quartz‐feldspar porphyry dyke 1* 0.001361 89 .047 12.0 .164 8.4 37 90 0.42 .007 2.7 0.5 43.7 1.2 .0262 74.9 0.02 75.0 .0066 3.5 .046 2 0.000407 37 .050 2.9 .057 2.7 169 1071 0.16 .014 0.6 13.3 92.3 0.5 .0435 6.2 0.09 6.3 .0143 0.6 .100 3 0.000078 96 .048 2.3 .053 2.9 173 1109 0.16 .014 0.6 13.7 92.0 0.5 .0467 3.3 0.09 3.4 .0144 0.6 .177 4 0.000215 53 .049 2.3 .054 2.9 177 1087 0.17 .014 0.6 13.2 90.5 0.5 .0462 4.5 0.09 4.5 .0141 0.6 .139 5 0.000052 211 .047 3.6 .181 2.5 198 377 0.54 .014 0.9 4.6 90.5 0.8 .0466 5.0 0.09 5.1 .0141 0.9 .180 6 0.001839 40 .077 12.3 .464 5.6 426 422 1.04 .015 1.0 5.5 93.1 1.5 .0505 29.8 0.10 29.8 .0146 1.7 .058 7* 0.000596 51 .053 4.1 .211 2.8 344 546 0.65 .007 1.0 3.5 47.0 0.5 .0440 11.5 0.04 11.6 .0073 1.2 .101 91 .050 2.1 .054 2.7 174 1144 0.16 .014 0.6 14.2 91.9 0.5 .0518 3.6 0.10 3.7 .0144 0.6 .159 0.000259 46 .050 1.8 .077 2.3 342 1586 0.22 .015 0.5 20.3 95.2 0.5 .0459 4.3 0.09 4.4 .0148 0.6 .132 2 .047 3.9 .150 2.9 149 347 0.44 .014 1.0 4.2 89.5 0.9 .0525 3.5 0.10 3.7 .0141 1.0 .266 11 0.000077 40 .049 2.4 .049 3.2 98 744 0.14 .014 0.6 8.9 89.0 0.6 .0480 2.7 0.09 2.8 .0139 0.6 .230 12 0.001341 45 .046 6.0 .129 4.8 52 149 0.36 .014 1.5 1.8 91.0 1.4 .0257 38.1 0.05 38.1 .0138 1.8 .048 8 9* 10 LW‐066 ‐ quartz porphyritic rhyolite breccias 1 1 .128 2.7 .218 2.2 109 545 0.44 .195 0.7 43.1 1079.8 12.5 .1276 2.7 3.44 2.7 .1953 0.7 .238 0.003032 31 .091 3.5 .223 3.4 74 1484 0.33 .016 1.2 3.2 99.6 1.4 .0464 35.0 0.10 35.1 .0155 2.2 .062 0.000481 48 .055 3.9 .143 3.7 127 710 0.40 .016 1.1 4.4 99.8 1.1 .0482 8.7 0.10 8.7 .0156 1.1 .131 4 .045 10.0 .182 7.4 142 4113 0.56 .004 2.2 0.8 24.1 0.5 .0561 8.4 0.03 8.7 .0038 2.2 .251 31 .050 6.8 .213 5.2 416 264 0.60 .002 1.6 1.4 14.4 0.2 .0147 86.7 0.00 86.8 .0022 2.1 .024 2 3 4 5 0.002298 6 0.000029 7 67 .115 1.2 .607 1.1 149 230 2.09 .342 1.0 21.7 1899.8 20.3 .1150 1.2 5.43 1.6 .3421 1.0 .648 51 .049 10.4 .150 8.9 241 328 0.46 .002 2.3 0.8 10.4 0.2 .0962 24.8 0.02 25.1 .0017 3.7 .149 8 0.001324 24 .065 2.5 .157 3.1 1487 29 0.37 .002 0.7 8.2 14.6 0.1 .0456 12.8 0.01 12.8 .0023 0.9 .072 9 0.000013 606 .097 1.5 .041 3.5 9 257 0.14 .273 1.1 15.5 1554.4 16.8 .0969 1.9 3.65 2.2 .2729 1.1 .501 10 0.001595 25 .095 2.7 .166 3.1 12 66 0.42 .193 1.6 4.8 1114.5 18.4 11 0.004730 22 .089 3.7 .283 3.2 739 74 0.51 .002 1.1 2.7 13.0 0.2 .0722 10.0 1.87 10.1 .1876 1.8 .0019 2.4 .179 Appendix A Page 214 Table A2 continued. Spot Name Meas 204
Pb/ 206
Pb Meas 207
Pb/ % 206
Pb err % err Meas 208
Pb/ 206
Pb Th
/238U Corr 206
Pb
/238U % err Th ppm U ppm % err 172 523 0.84 .100 0.7 18.2 232
Rad 206
Pb (ppm) 207 corr 206
Pb /238U Age 204 corr 207
Pb /206Pb 1σ err 204 corr 207
Pb
/235U % err % err 204 corr 206Pb
/238U % err err corr LW‐026 1 0.000175 71 .063 2.6 .265 2.1 2 0.000043 35 .073 1.1 .002 9.5 3 495 0.01 .158 0.7 55.0 942.7 6.1 .0727 1.2 1.58 1.4 .1579 0.7 .483 3 0.000016 67 .161 4.7 .103 1.4 47 207 0.32 .454 1.1 59.4 2395.7 41.7 .1609 4.7 10.08 4.8 .4543 1.1 .229 4 0.000085 41 .050 2.6 .251 1.8 334 716 0.76 .030 0.7 11.7 191.6 1.4 .0487 2.9 0.20 3.0 .0301 0.7 .247 5 0.000016 3231 .052 10.0 .141 7.2 191 452 0.40 .004 1.6 1.7 25.8 0.4 .0515 18.0 0.03 18.1 .0040 1.8 .102 6 0.000140 48 .051 2.2 .234 1.6 292 393 0.77 .045 0.7 15.3 284.7 1.9 .0494 3.1 0.31 3.2 .0450 0.7 .214 0.000094 41 .071 1.2 .100 1.5 95 211 0.32 .151 0.6 39.9 906.6 5.1 .0692 1.4 1.44 1.6 .1510 0.6 .373 4 .050 7.9 .139 7.7 83 307 0.41 .007 1.9 1.3 47.2 0.9 .0677 6.8 0.07 7.0 .0075 1.9 .265 7 8 614.4 4.2 .0607 4.1 0.84 4.1 .1001 0.7 .172 9 0.001478 28 .052 3.1 .177 2.6 349 405 0.50 .014 0.8 8.9 92.2 0.8 .0294 24.6 0.06 24.6 .0141 1.1 .045 10 0.000047 46 .080 0.9 .124 1.1 115 294 0.40 .196 0.6 49.4 1149.6 6.3 .0792 1.0 2.13 1.2 .1955 0.6 .472 11 4 .052 9.5 .175 7.9 241 152 0.48 .002 2.2 0.8 11.8 0.3 .0958 8.0 0.03 8.3 .0019 2.2 .266 LW‐033 ‐quartz‐feldspar porphyritic rhyolite dyke 1.1 0.000267 47 .047 3.8 .075 4.6 364 2097 0.23 .003 0.9 4.9 21.8 0.2 .0433 6.1 0.02 6.2 .0034 0.9 .148 1.1 0.000302 17 .183 0.6 .142 1.9 94 475 0.57 .485 0.7 70.7 2490.6 24.2 .1794 0.8 11.94 1.1 .4827 0.7 .676 2 0.000082 52 .082 1.1 .139 1.3 86 1671 0.45 .199 0.7 33.9 1167.5 7.6 .0804 1.4 2.21 1.6 .1989 0.7 .431 4 0.000883 43 .052 5.1 .126 5.1 727 98 0.36 .001 1.3 2.5 9.0 0.1 .0393 16.7 0.01 16.8 .0014 1.4 .086 5 0.005973 32 .086 9.2 .183 5.9 30 3219 0.32 .016 1.9 1.4 98.2 2.2 .0144 4.3 6 0.033547 8 .480 3.1 1.264 3.9 365 1689 0.79 .008 1.2 3.1 22.8 3.3 .0030 15.6 7 0.000056 49 .083 0.7 .091 1.2 154 1533 0.30 .218 0.4 99.6 1271.5 5.3 .0820 0.8 2.46 1.0 .2177 0.4 .447 10 0.000037 78 .051 1.1 .020 2.7 94 156 0.06 .037 0.3 48.7 234.0 0.8 .0505 1.4 0.26 1.5 .0369 0.3 .235 11 0.003089 5 .093 2.9 .122 6.1 314 98 0.10 .030 0.3 82.7 179.9 1.5 .0479 15.7 0.19 15.8 .0282 0.5 .033 12 0.000089 16 .052 1.2 .018 3.1 82 198 0.05 .036 0.3 51.8 225.8 0.8 .0507 1.3 0.25 1.4 .0357 0.3 .250 0.000091 35 .129 0.7 .192 0.9 132 532 0.57 .252 0.5 52.3 1378.0 12.1 .1280 0.8 4.44 1.0 .2513 0.5 .551 1 .060 2.5 .053 4.0 15 242 0.16 .097 1.0 8.1 596.5 5.9 .0618 2.5 0.83 2.7 .0972 1.0 .372 86 .057 2.5 .206 2.0 99 170 0.66 .066 0.9 8.8 410.3 3.6 .0554 3.9 0.50 4.0 .0658 0.9 .223 13 14 15 0.000128 Appendix A Page 215 Figure A1. SEM Cathodoluminesence images of zircons showing location of spot analyses for SHRIMP‐RG data. Appendix A Page 216 A.2 Additional U‐Pb Zircon Geochronologic Data Sensitive High Resolution Ion Microprobe ‐ Reverse Geometry (SHRIMP‐RG) data are provided here for 4 samples not reported elsewhere in the thesis. Two samples, LW‐30 and LW‐31, were collected by the author in the Amotape range to the west of the main study area (refer to Fig. A2; locations and sample descriptions are given in table A3). Samples 1601 and 3763 were provided to the author by personnel working for BHP Billiton and represent mineralized porphyry Cu type occurrences within the Lancones basin. Sample 3763, the Chancadora porphyry prospect, is within the study area whereas 1601, the Linderos porphyry prospect, is located in Ecuador. No detailed description was available for these units, though the Linderos porphyry prospect has been referred to with respect to a lead isotope study by Chiaradia et al. (2004). A.3 Results Sample 3763, a granodiorite porphyry, yielded a weighted mean 206Pb/238U age of 78.7 ± 1.4 Ma (MSWD=3.2), based on 11 analyses and one rejected analyses (Fig. A3). Sample 1601, also a granodiorite porphyry, yield a weighted mean 206Pb/238U age of 88.1 ± 1.4 Ma (MSWD=8.5) based on 11 single grain analyses (Fig. A4). The 78.7 Ma age determined for sample 3763 is broadly similar to a few of the Ar‐Ar ages determined for granitic rocks in the regions (Appendix B). However, the older age determined for sample 1601 is slightly younger than the youngest of the volcanic rocks dated in the Lancones basin (Chap. 2). Samples from the Amotape range represent basement rocks of the northwest Peruvian margin. Though most of the 19 analyses of zircons from sample LW‐30 yielded a highly scattered range of 206Pb/238U ages representing Paleozoic and Proterozoic inheritance, 8 analyses form a relatively tight cluster and give a weighted mean 206Pb/238U age of 230.2 ± 2.9 Ma (MSWD=65; Fig. A5). This age is in very good agreement with Middle to Late Triassic U‐Pb ages reported from the northernmost part of the Amotape range in Ecuador (Noble et al., 1997). Sample LW‐31 yields an abundance of Proterozoic inherited zircons, though three analyses cluster ca. 230 Ma (Fig. A6). A few samples display probable Pb loss and no age has been calculated for this sample. Appendix A Page 217 Table A3. Description of additional rock samples for U‐Pb zircon analysis (not included in the Chapters). Sample ID UTM Easting UTM Northing 3763 556970 9474559 1601 601176 9528120 LW‐030 504492 9430344 LW‐031 484306 9426178 Appendix A Calculated age +/‐ error (Ma) Description
78.6 +/‐ 1.9 "Chancadora Porphyry" prospect. Granodioritic porphyry with "A" veining. 87.5 +/‐1.6 "Linderos Porphyry" prospect. Granodioritic porphyry with veining. 229.8 +/‐ Biotite +/‐hornblende‐bearing 2.9 coarse grained meta‐granodiorite w/ some xenoliths. not Muscovite‐biotite‐bearing med‐
determined grained granite Analytical method Comments
SHRIMP
Pyrite floated off; a few very fine grained zircons SHRIMP
abundant zircons.
SHRIMP
abundant zircons
SHRIMP
abundant zircons
Page 218 Table A4. U‐Pb zircon data from SHRIMP‐RG analysis. Spot Name Meas 204
Pb/ 206
Pb % err Meas 207
Pb/ 206
Pb % err Meas 208
Pb/ 206
Pb % err Th ppm U ppm 232
Th
/238U Corr 206
Pb
/238U % err Rad 206
Pb (ppm) .012 2.8 0.7 207 corr 206
Pb /238U Age 1σ err 204 corr 207
Pb /206Pb 204 corr 207
Pb /235U % err .0528 8.4 % err 204 corr 206Pb
/238U % err err corr 8.9 .0124 2.8 .315 .059 .119 .062 .105 .102 .168 .134 .054 .128 .118 3763 ‐ "Chancadora Porphyry" prospect. Granodioritic porphyry with "A" veining. 1 0.000000 0 .053 8.4 .124 7.8 2 0.001592 35 .052 6.7 .213 3 0.000898 40 .051 8.3 .137 4 0.001547 41 .052 7.7 .216 5.2 5 0.000657 33 .053 3.2 .115 2.9 6 0.001168 43 .057 6.4 .127 5.9 7 0.000359 41 .050 4.7 .123 5.3 8 ‐0.000647 54 .051 4.1 .082 9 0.001194 62 .047 7.0 .102 10 0.000778 21 .050 4.4 .094 11 0.000005 7280 .050 4.6 .180 0.31 79.1 2.2 0.09 19 62 4.5 66 114 0.60 .013 1.7 1.2 79.8 1.4 .0282 34.0 0.05 34.1 .0122 2.0 6.8 24 76 0.33 .012 2.1 0.8 78.0 1.7 .0378 18.5 0.06 18.7 .0120 2.2 46 83 0.57 .012 2.0 0.9 78.5 1.6 .0289 37.5 0.05 37.6 .0120 2.3 241 604 0.41 .012 0.8 6.2 76.2 0.6 .0430 0.07 .0118 0.9 36 112 0.33 .013 2.0 1.2 81.3 1.7 .0391 21.8 0.07 21.9 .0126 2.2 84 238 0.36 .013 1.2 2.6 80.7 1.0 .0448 7.2 0.08 7.3 .0126 1.2 4.2 69 315 0.23 .013 1.1 3.4 80.6 0.9 .0604 9.1 0.11 9.1 .0128 1.2 9.0 30 112 0.27 .013 1.7 1.2 80.7 1.4 .0293 40.9 0.05 41.0 .0123 2.2 4.3 70 271 0.27 .012 1.1 2.9 78.4 0.9 .0381 0.06 9.1 .0121 1.2 3.2 134 258 0.54 .013 1.2 2.9 83.6 1.0 .0500 11.1 0.09 11.2 .0131 1.3 8.6 9.0 8.6 1601 ‐ "Linderos Porphyry" prospect. Granodioritic porphyry with veining. 1 0.002663 48 .054 9.2 .079 10.2 11 52 0.22 .014 2.5 0.6 86.1 2.2 .0129 3.5 2 0.001124 67 .053 7.9 .083 8.4 13 71 0.19 .014 2.1 0.9 90.5 2.0 .0362 33.9 0.07 34.0 .0139 2.5 3 0.000000 0 .054 6.0 .076 6.6 24 125 0.20 .014 1.6 1.5 87.2 1.4 .0543 6.0 0.10 6.3 .0137 1.6 4 ‐0.000769 3 .048 6.2 .066 6.8 26 138 0.19 .014 1.5 1.6 87.9 1.4 .0594 5.1 0.11 5.3 .0139 1.5 5 0.001286 66 .048 8.4 .101 7.9 16 69 0.24 .013 2.1 0.8 84.4 1.8 .0291 47.4 0.05 47.5 .0129 2.6 6 0.000946 72 .048 6.7 .076 7.0 25 113 0.23 .014 1.7 1.4 91.8 1.6 .0336 32.4 0.07 32.4 .0141 2.1 7 0.000000 0 .057 7.3 .073 8.9 16 80 0.20 .014 2.0 0.9 85.9 1.8 .0572 0.11 7.6 .0136 2.0 8 0.001973 26 .049 8.0 .093 7.8 18 72 0.26 .014 2.1 0.9 91.2 1.9 .0188 48.8 0.04 48.9 .0138 2.3 9 0.000717 41 .050 4.0 .214 5.8 173 279 0.64 .014 1.0 3.4 89.4 0.9 .0390 12.6 0.07 12.6 .0138 1.2 10 ‐0.000381 50 .051 5.0 .093 5.0 51 187 0.28 .014 1.3 2.2 87.4 1.2 .0570 0.11 6.8 .0138 1.3 11 0.001837 41 .047 5.2 .083 5.7 37 172 0.22 .013 1.3 2.0 86.3 1.2 .0190 64.4 0.03 64.5 .0130 1.9 7.3 6.6 Appendix A Page 219 .074 .258 .285 .056 .064 .265 .047 .093 .199 .030 Table A4 continued. Spot Name Meas 204
Pb/ 206
Pb % err Meas 207
Pb/ 206
Pb % err Meas 208
Pb/ 206
Pb % err Th ppm U ppm 232
Th
/238U Corr 206
Pb
/238U % err Rad 206
Pb (ppm) 207 corr 206
Pb /238U Age 1σ err 204 corr 207
Pb /206Pb 204 corr 207
Pb /235U % err .0515 1.9 % err 204 corr 206Pb
/238U % err err corr 2.0 .0368 0.4 .189 .338 .313 .258 .147 .374 .262 .242 .225 .279 .329 .189 .372 .417 .339 .343 .368 .524 .234 LW‐30 ‐ Biotite +/‐hornblende‐bearing coarse grained meta‐granodiorite w/ some xenoliths 1 0.000151 32 .054 1.2 .024 2.9 81 1441 2 0.000184 51 .069 2.5 .221 2.2 45 1520 3 0.000024 41 .050 1.2 .018 3.0 78 2332 4 0.000068 22 .051 1.1 .018 2.7 177 1866 5 0.000217 22 .051 1.0 .007 4.2 54 1737 6 0.000034 39 .062 1.3 .024 3.2 26 587 0.06 .037 0.4 45.7 232.8 0.68 .134 0.05 .037 0.08 0.03 0.08 0.9 0.26 1.2 7.9 809.3 9.5 .0667 3.4 1.23 3.6 .1339 1.2 0.4 48.6 235.5 1.0 .0499 1.3 0.26 1.3 .0372 0.4 .036 0.3 71.8 227.1 0.8 .0497 1.2 0.25 1.3 .0358 0.3 .036 0.3 58.1 229.3 0.7 .0479 2.1 0.24 2.1 .0361 0.3 .103 0.6 28.8 628.8 3.5 .0614 1.4 0.87 1.5 .1026 0.6 7 0.000001 2078 .052 1.1 .015 3.0 80 1324 0.05 .036 0.3 53.5 226.7 0.7 .0516 1.2 0.25 1.2 .0358 0.3 8 0.000152 21 .051 1.9 .058 2.7 103 528 0.18 .036 0.6 18.2 228.3 1.3 .0490 2.3 0.24 2.4 .0360 0.6 9 0.000090 38 .051 1.3 .019 3.1 88 366 0.07 .037 0.4 42.3 235.4 0.9 .0493 1.7 0.25 1.7 .0371 0.4 20 0.000085 14 .051 1.9 .020 4.6 32 305 0.06 .036 0.6 16.4 228.8 1.3 .0496 2.0 0.25 2.1 .0361 0.6 21 0.000107 42 .054 1.5 .056 2.7 50 210 0.14 .055 0.7 17.4 347.1 2.4 .0528 2.0 0.40 2.2 .0553 0.7 22 0.000029 34 .069 2.3 .085 2.9 161 88 0.25 .128 0.4 73.4 771.3 3.7 .0685 2.3 1.21 2.4 .1277 0.4 23 0.000096 22 .061 1.4 .165 1.6 122 212 0.41 .080 0.6 21.0 494.0 3.1 .0593 1.6 0.65 1.7 .0799 0.6 24 ‐0.000135 56 .077 1.5 .269 1.6 83 340 0.87 .181 0.9 15.5 1069.9 9.7 .0791 2.0 1.98 2.2 .1815 0.9 25 0.000159 39 .059 2.6 1.342 1.2 351 326 4.13 .091 1.1 6.9 560.8 6.3 .0568 3.2 0.71 3.4 .0907 1.1 26 0.000038 67 .058 1.9 .444 1.1 286 668 1.39 .097 0.7 17.6 596.1 4.3 .0572 2.0 0.76 2.1 .0966 0.7 27 ‐0.000058 59 .057 1.8 .218 1.7 138 69 0.68 .080 0.8 14.4 496.2 3.9 .0581 2.0 0.64 2.1 .0801 0.8 28 0.000082 52 .080 1.3 .135 1.8 45 120 0.39 .178 1.0 18.4 1050.5 10.0 .0790 1.6 1.94 1.9 .1780 1.0 29 0.000051 44 .063 2.6 .128 3.8 133 99 0.40 .098 0.6 28.5 597.8 3.9 .0622 2.7 0.84 2.7 .0975 0.6 Appendix A Page 220 Table A4 continued. Spot Name Meas 204
Pb/ 206
Pb % err Meas 207
Pb/ 206
Pb % err Meas 208
Pb/ 206
Pb % err Th ppm U ppm 232
/
207 corr 206
Pb /238U Age Th
U Corr 206
Pb
/238U % err Rad 206
Pb (ppm) 110.7 86.8 2.0 238
1σ err 204 corr 207
Pb /206Pb 204 corr 207
Pb /235U % err 204 corr 206Pb
/238U % err err corr .0489 44.7 0.09 44.7 .0136 1.4 .032 .025 .049 .037 .162 .320 .129 .565 .506 .219 .685 .350 .397 .608 .348 .398 .334 .186 .656 .405 .467 .695 .481 .505 .305 .380 .349 .616 .426 % err LW‐31 ‐ Muscovite‐biotite‐bearing med‐grained granite 1 0.008481 6 .174 2.0 .326 2.7 239 1632 0.03 .016 0.3 2 3 0.008222 5 .160 0.7 .291 1.5 164 8038 0.04 .029 0.4 92.7 157.4 3.1 .0371 55.7 0.12 55.7 .0244 1.4 0.002552 10 .086 1.4 .096 3.7 154 5149 0.03 .025 0.4 112.8 155.0 1.1 .0484 14.0 0.16 14.1 .0243 0.7 4 5 0.002809 5 .093 1.1 .116 2.2 356 3772 0.10 .027 0.2 89.5 162.7 1.2 .0522 12.0 0.18 12.0 .0256 0.4 0.000164 38 .051 1.3 .003 7.3 18 3856 0.01 .036 0.4 60.5 227.2 0.9 .0487 2.4 0.24 2.4 .0358 0.4 6 10 ‐0.000243 1 .053 2.9 .215 2.2 83 117 0.68 .067 1.0 7.3 419.5 4.0 .0566 2.9 0.53 3.0 .0673 1.0 0.000949 20 .052 3.7 .054 6.4 19 1964 0.16 .036 1.3 3.6 228.7 3.0 .0375 10.4 0.18 10.5 .0355 1.4 11 12 ‐0.000015 0 .072 0.5 .111 1.3 341 399 0.36 .166 0.4 137.9 987.5 3.4 .0721 0.5 1.65 0.6 .1656 0.4 0.000044 35 .075 0.5 .131 0.9 336 227 0.37 .164 0.4 132.6 972.9 3.5 .0747 0.6 1.68 0.7 .1636 0.4 13 0.000162 48 .050 2.0 .065 3.2 79 471 0.21 .036 0.7 12.4 228.7 1.7 .0476 3.3 0.24 3.4 .0360 0.7 14 0.000032 40 .114 0.7 .124 2.4 100 1052 0.44 .323 0.6 65.6 1794.4 11.6 .1133 0.7 5.04 0.9 .3225 0.6 15 0.000059 17 .053 1.5 .146 1.6 212 126 0.47 .057 0.6 23.0 357.2 2.0 .0520 1.5 0.41 1.6 .0569 0.6 16 0.000012 30 .101 0.4 .013 2.2 153 236 0.05 .205 0.2 593.3 1172.6 4.7 .1007 0.4 2.85 0.5 .2052 0.2 17 0.000042 35 .099 0.7 .135 1.1 109 75 0.44 .284 0.6 62.0 1611.8 8.9 .0982 0.7 3.84 0.9 .2836 0.6 18 0.000025 45 .053 0.9 .100 1.2 323 571 0.32 .060 0.4 54.5 377.7 1.4 .0528 1.0 0.44 1.1 .0602 0.4 19 0.000044 66 .061 1.7 .138 2.0 73 180 0.42 .101 0.8 15.5 618.1 4.9 .0605 1.8 0.84 2.0 .1006 0.8 20 0.000165 40 .056 2.9 .272 2.4 61 666 0.84 .086 1.3 5.5 532.9 6.6 .0534 3.6 0.63 3.8 .0857 1.3 21 0.000208 66 .047 4.0 .188 3.7 975 432 0.62 .002 1.2 3.5 15.9 0.2 .0437 6.4 0.01 6.5 .0025 1.2 22 0.000071 40 .112 1.4 .545 1.2 82 289 1.79 .331 1.3 13.4 1847.7 23.9 .1108 1.5 5.06 1.9 .3310 1.3 23 0.000034 27 .059 1.0 .187 1.0 327 943 0.59 .094 0.5 46.4 582.6 2.6 .0581 1.0 0.76 1.1 .0944 0.5 24 0.000076 36 .073 1.2 .254 1.2 133 168 0.82 .164 0.7 23.7 980.6 7.1 .0716 1.4 1.62 1.6 .1642 0.7 25 0.000000 0 .177 1.4 .139 2.1 137 970 0.49 .428 1.4 107.1 2190.5 37.8 .1774 1.4 10.48 2.0 .4285 1.4 26 0.000035 45 .073 1.1 .032 7.5 29 3366 0.11 .148 0.6 36.7 883.6 5.5 .0727 1.1 1.48 1.3 .1477 0.6 27 ‐0.000139 55 .096 1.2 .205 1.5 55 86 0.67 .263 0.9 19.4 1500.9 14.1 .0976 1.6 3.55 1.9 .2635 1.0 28 0.000000 0 .060 1.4 .100 4.6 189 254 0.29 .110 0.5 63.1 675.7 3.1 .0603 1.4 0.92 1.5 .1103 0.5 29 0.000037 67 .058 1.6 .626 1.0 451 237 1.97 .085 0.7 17.3 527.1 3.7 .0576 1.8 0.68 1.9 .0852 0.7 30 ‐0.000164 1 .052 2.5 .207 2.3 144 47 0.66 .042 0.9 8.1 263.6 2.4 .0548 2.4 0.32 2.6 .0419 0.9 31 0.000032 16 .163 1.5 .102 1.0 65 291 0.36 .454 1.2 72.5 2381.7 33.6 .1631 1.5 10.19 1.9 .4533 1.2 32 0.000041 5 .064 1.0 .121 1.3 148 186 0.35 .119 0.5 44.1 723.7 3.5 .0629 1.0 1.03 1.2 .1187 0.5 Appendix A Page 221 Figure A2. Regional geologic map of northwestern Perú showing the location of U‐Pb zircon geochronology samples. Appendix A Page 222 box heights are 2σ
110
Age = 78.7 ± 1.4
(MSWD = 3.2)
90
206
Pb/238U age (Ma)
100
80
70
Figure A3. Box plot for all sample points for sample LW‐30 illustrating 207Pb‐corrected 206
Pb*/238U data with error bars at 2σ. Open boxes are omitted whereas solid boxes were included in the age calculation. Ages given are 206Pb/238U with 2σ uncertainties. box heights are 2σ
110
Age = 88.1 ± 1.4
(MSWD = 8.5)
90
206
Pb/238U age (Ma)
100
80
70
Figure A4. Box plot for all sample points for sample LW‐31 illustrating 207Pb‐corrected 206
Pb*/238U data with error bars at 2σ. Open boxes are omitted whereas solid boxes were included in the age calculation. Ages given are 206Pb/238U with 2σ uncertainties. Appendix A Page 223 1200
box heights are 2σ
800
600
206
Pb/
238
U age (Ma)
1000
400
200
Age 230.2 ± 2.9
(MSWD = 65)
0
Figure A5. Box plot for all sample points for sample 3763 illustrating 207Pb‐corrected 206
Pb*/238U data with error bars at 2σ. Open boxes are omitted whereas solid boxes were included in the age calculation (and are noted by those encircled with the dashed line). Ages given are 206Pb/238U with 2σ uncertainties. 2500
box heights are 2σ
1500
1000
206
Pb/238U age (Ma)
2000
500
0
Figure A6. Box plot for all sample points for sample 1601 illustrating 207Pb‐corrected 206
Pb*/238U data with error bars at 2σ. No age was determined for this sample, though the circled data points represent possible igneous age of the sample. Ages given are 206Pb/238U with 2σ uncertainties. Appendix A Page 224 A.4 References Chiaradia, M., Fontbote, L., and Paladines, A. 2004. Metal sources in mineral deposits and crustal rocks of Ecuador (1 degrees N‐4 degrees S); a lead isotope synthesis, Economic Geology and the Bulletin of the Society of Economic Geologists, 99: 1085‐1106. Krogh, T.E. 1982. Improved accuracy of U‐Pb zircon ages by the creation of more concordant systems using an air abrasion technique. Geochimica et Cosmochimica Acta, 46: 637‐649. Noble, S.R., Aspden, J.A., and Jemielita, R.A. 1997. Northern Andean crustal evolution; new U‐
Pb geochronological constraints from Ecuador. Geological Society of America Bulletin, 109: 789‐798. Mortensen, J.K., Ghosh, D., and Ferri, F. 1995. U‐Pb age constraints of intrusive rocks associated with Copper‐Gold porphyry deposits in the Canadian Cordillera, in Schroeter, T.G., ed., Porphyry deposits of the northwestern Cordillera of North America: Canadian Institute of Mining and Metallurgy, Special Volume 46, pp. 142‐158. Roddick, J.C. 1987. Generalized numerical error analysis with application to geochronology and thermodynamics. Geochimica et Cosmochimica Acta, 51: 2129‐2135. Stacey, J.S. and Kramer, J.D. 1975. Approximation of terrestrial lead isotope evolution by a two‐stage model. Earth and Planetary Science Letters, 26: 207‐221. Steiger, R.H. and Jäger, E. 1977. Subcommission on geochronology: convention on the use of decay constants in geo‐ and cosmochronology. Earth and Planetary Science Letters, 36: 359‐362. Appendix A Page 225 Appendix B. Ar‐Ar Geochronologic Data B.1 Methodology Six samples were dated using the 40Ar‐39Ar analytical method at PCIGR‐UBC, including 5 samples of intrusive rocks from batholiths or dykes and 1 volcaniclastic rock sample. Sample descriptions and locations are provided in Table B1 and Figure B1. Each sample was crushed and sieved to obtain fragments typically less than 2 mm. A hand magnet was passed over the samples to remove magnetic minerals and metallic crusher fragments/spall. The samples were washed in deionized water, rinsed and then air‐dried at room temperature. Mineral separates were hand‐picked, wrapped in aluminum foil and stacked in an irradiation capsule with similar‐aged samples and neutron flux monitors (Fish Canyon Tuff sanidine, 28.02 Ma; Renne et al., 1998; MAC‐83 biotite, 24.36; Sandeman et al., 1999). The samples were irradiated on October 14 through 16, 2003 at the McMaster Nuclear Reactor in Hamilton, Ontario, for 72 MWH, with a neutron flux of approximately 3x1016 neutrons/cm2. Analyses (n=33) of 11 neutron flux monitors irradiated with the samples produced errors of <0.25% in the J value. The samples were analyzed on March 9 and 10, 2004, at the Noble Gas Laboratory, Pacific Centre for Isotopic and Geochemical Research, University of British Columbia, Vancouver, BC, Canada. The separates were step‐heated at incrementally higher powers in the defocused beam of a 10W CO2 laser (New Wave Research MIR10) until fused. The gas evolved from each step was analyzed by a VG5400 mass spectrometer equipped with an ion‐counting electron multiplier. All measurements were corrected for total system blank, mass spectrometer sensitivity, mass discrimination, radioactive decay during and subsequent to irradiation, as well Appendix B page 226 as interfering Ar from atmospheric contamination and the irradiation of Ca, Cl and K. Details of the analyses are presented in Table B2. Errors are quoted at the 2‐sigma (95% confidence) level and are propagated from all sources except mass spectrometer sensitivity and age of the flux monitor. B.2 Results All analyses reported were performed on hornblende mineral separates. Several feldspar and biotite mineral separates were analyzed but yielded more variable results. Sample LW‐06, a hornblende granodiorite pluton exposed for several kilometres in the western sector of the map area, yielded a plateau age of 47.1 ± 3.1 Ma with MSWD of 0.85 using 100% of the 39Ar (Fig. B2). Located close to LW‐06, sample LW‐07 was collected from a volcaniclastic unit and yielded a plateau age of 48.6 ± 7.4 Ma with MSWD of 1.3 using 75% of the 39Ar (Fig. B3). The age is interpreted as a reset due to the thermal effects of the proximal granodiorite intrusion. Sample LW‐36 (hornblende diorite) is located approximately 10 km to the east of samples LW‐06 and ‐07, and very similar in composition as LW‐06. This sample produced a plateau age of 52.2 Ma ± 3.4 Ma with MSWD of 0.99 including 96% of the 39Ar (Fig. B4). Sample LW‐88, a hornblende porphyry dyke northeast of the Los Lomas batholith, yielded a similar but less precise plateau age of 55.0 Ma ± 3.8 Ma with MSWD of 1.7 factoring in 44% of the 39Ar (Fig. B5). Sample LW‐85 is a hornblende granite of the Los Lomas batholith and yields a plateau age of 72.4 Ma ± 3.4 Ma with MSWD of 1.4 including 95.5% of the 39Ar (Fig. B6). A similar age was determined from LW‐61, a hornblende and plagiclase porphyritic dyke southeast of the batholith. LW‐61 yielded a plateau age of 74.9 Ma ± 3.3 Ma with MSWD of 1.4 including 97.3% of the 39Ar (Fig. 7). Appendix B page 227 Table B1. 40Ar‐39Ar rock sample description and location data. Eastings and northings are UTM Zone 17, Southern Hemisphere (WGS84). Sample
LW-06
LW-07
LW-36
LW-88
LW-85
LW-61
Rock Type
hornblende granodiorite
hornblende-plagioclase +/quartz porphyritic rhyolite
(volcaniclastic)
hornblende-bearing diorite
hornblende porphyry dyke
hornblende-bearing granite
hornblende-plagioclase
porphyry dyke
Mineral
Analyzed
hornblende
Easting
559404
Northing
9481200
hornblende
hornblende
hornblende
hornblende
560758
567854
588812
581496
9483392
9479640
9490854
9485511
hornblende
584704
9466924
Appendix B page 228 1
Laser
Power
(%)
Isotope Ratios
40
Ar/39Ar
38
Ar/39Ar
37
Ar/39Ar
36
Ar/39Ar
39
%40Ar
Cl/K
Ca/K
-10
%39Ar
40
Ar*/39ArK
Age±2σ
3
LW-06 (granodiorite), hornblende, J = 0.008961±0.000016; volume ArK = 68.82 x 10 cm , integrated age = 48.06±5.81 Ma (2σ)
0.524±0.088
0.555±0.041
0.622±0.066
1.264
0.093
101.71
1.42-2.913±9.923
-47.72±164.71
2.553.563 0.020
0.186 0.038
0.118 0.024
0.169 0.066
0.477
0.032
92.38
6.953.852 3.183
61.22 49.74
320.973 0.014
0.512 0.018
0.127 0.020
0.060 0.074
0.563
0.113
79.71
6.53.650 1.316
58.07 20.60
3.712.038 0.010
2.298 0.014
0.027 0.637
0.031 0.046
0.552
0.525
72.47
41.123.154 0.423
50.28 6.64
3.710.434 0.014
2.338 0.019
0.054 0.326
0.027 0.031
2.106
0.535
70.56
33.012.873 0.234
45.86 3.69
3.942.628 0.023
2.022 0.042
0.307 0.023
0.135 0.069
0.773
0.462
88.07
2.594.238 2.680
67.24 41.74
4.227.371 0.046
2.613 0.047
0.222 0.052
0.088 0.106
2.74
0.599
89.68
4.212.351 2.589
37.61 41.00
5.528.936 0.024
3.496 0.028
0.202 0.039
0.094 0.065
1.328
0.803
90.66
4.212.271 1.715
36.34 27.17
2179.290±0.041
39
-10
3
LW-07 (rhyolitic volcaniclastic), hornblende, J = 0.008959±0.000016; volume ArK = 59.67 x 10 cm , integrated age = 30.29±7.66 Ma (2σ)
284.108±0.033
0.116±0.092
0.328±0.048
0.281±0.062
1.095
0.012
96.99
2.842.298±4.463
2.357.582 0.032
0.085 0.117
0.218 0.399
0.204 0.065
2.431
0.008
103.62
4.86-1.962 3.635
-31.99 59.80
2.655.344 0.049
0.168 0.072
0.333 0.390
0.223 0.087
5.601
0.026
119.3
3.49-9.606 4.926
-162.34 87.11
2.956.241 0.062
0.673 0.080
0.534 0.112
0.222 0.104
5.737
0.145
115.88
1.94-7.331 5.689
-122.56 98.42
3.214.654 0.026
2.293 0.033
0.092 0.051
0.045 0.057
1.124
0.524
84.53
11.81.955 0.693
31.33 11.01
3.414.165 0.023
2.610 0.033
0.088 0.172
0.040 0.097
1.822
0.598
76.04
14.283.004 1.114
47.91 17.54
3.610.680 0.021
2.418 0.029
0.058 0.143
0.025 0.109
1.208
0.553
62.39
21.833.613 0.814
57.47 12.74
3.810.875 0.019
2.463 0.019
0.058 0.133
0.028 0.099
1.175
0.564
67.92
21.853.141 0.806
50.06 12.67
420.564 0.033
2.652 0.033
0.164 0.108
0.066 0.093
2.585
0.608
89.64
7.051.790 1.756
28.70 27.93
4.334.468 0.030
3.272 0.035
0.305 0.097
0.119 0.101
2.353
0.753
96.66
3.260.908 3.429
14.62 54.97
36.76±70.68
4.826.172 0.038
3.663 0.041
0.217 0.070
0.088 0.091
1.868
0.843
93
4.681.489 2.168
23.91 34.58
645.223 0.057
3.231 0.061
0.447 0.060
0.155 0.109
2.33
0.745
95.18
2.131.689 4.356
27.10 69.36
Table B2. 40Ar‐39Ar age data for plutonic and volcanic rock samples from the Lancones basin. Neutron flux monitors: 24.36 Ma MAC‐83 biotite (Sandeman et al. 1999); 28.02 Ma FCs (Renne et al., 1998). Isotope production ratios: (40Ar/39Ar)K=0.0302, (37Ar/39Ar)Ca=1416.4306, (36Ar/39Ar)Ca=0.3952, Ca/K=1.83(37ArCa/39ArK).
Appendix B page 229 1
Laser
Isotope Ratios
Power
(%)
40
38
37
36
40
Ar/39Ar
Ar/39Ar
Ar/39Ar
Ar/39Ar
Ca/K Cl/K %40Ar %39Ar
Ar*/39ArK
Age±2σ
39
-10
3
LW-36 (diorite), hornblende, J = 0.008957±0.000014; volume ArK = 119.09 x 10 cm , integrated age = 45.60±5.34 Ma (2σ)
2807.088±0.133
4.622±0.148
4.732±0.133
3.176±0.159
18.431
1.1
116.3
0.08-134.518±97.834
2.3420.678 0.295
1.080 0.295
1.458 0.301
1.792 0.621 101.78
0.41-7.308 28.118
2.883 0.300
-122.14 486.20
0.363 0.061
0.454 0.087
1.624 0.164
103
1.3-3.700 8.991
2.6129.388 0.058
0.801 0.065
-60.81 150.27
2.988.059 0.038
0.210 0.041
0.308 0.057
0.835 0.104 102.88
2.25-2.439 4.031
0.517 0.045
-39.85 66.59
0.239 0.037
0.239 0.076
1.737 0.073
97.55
2.081.601 4.878
3.271.519 0.033
0.372 0.063
25.69 77.72
0.113 0.038
0.125 0.062
0.752 0.075
93.52
4.382.337 2.003
3.538.914 0.031
0.363 0.049
37.37 31.71
0.
0.0
0.
3
80.8
29.6
3.818.493
0.011
1.742
0.014
026
0.854
52
0.022
785
0.
96
5
3.436
0.313
54.69 4.91
0.035 0.046
0.052 0.051
0.278 0.326
79.64
418.810 0.021
1.437 0.025
14.633.645 0.738
57.95 11.54
0.041 0.123
0.033 0.078
1.261 0.456
73.35
4.212.455 0.014
1.998 0.023
19.483.138 0.742
50.01 11.66
4.412.268 0.015
2.086 0.022
0.061 0.073
0.033 0.050
1.342 0.477
74.52
10.712.837 0.479
45.27 7.54
52.82 11.66
4.813.818 0.017
2.141 0.018
0.051 0.025
0.037 0.068
0.279 0.489
73.6
9.543.317 0.743
5.516.820 0.023
0.105 0.055
0.048 0.070
1.678 0.509
77.22
5.533.360 0.981
2.228 0.025
53.50 15.40
39
-10
3
LW-88 (hornblende porphyritic dyke), hornblende, J = 0.008948±0.000016; volume ArK = 235.18 x 10 cm , integrated age = 56.86±2.77 Ma (2σ)
24172.030±0.287
3.098±0.304
5.441±0.288
15.269±0.290
41.196 0.085 107.03
0.01-475.602±321.562
2.3351.623 0.046
0.352 0.066
0.676 0.054
1.262 0.074
4.794 0.024 103.94
-225.07 418.93
0.18-13.108 22.908
2.683.108 0.039
0.087 0.129
0.160 0.061
0.276 0.077
1.346 0.005
58.74 90.68
95.06
0.793.699 5.803
92.22
1.684.052 2.640
2.956.227 0.016
0.073 0.059
0.085 0.087
0.180 0.049
1.308 0.006
64.26 41.12
94.42
2.322.107 2.891
3.241.108 0.031
0.061 0.108
0.063 0.063
0.134 0.079
1.022 0.005
33.69 45.80
3.430.804 0.013
97.04
2.430.800 2.104
0.057 0.067
0.065 0.092
0.104 0.069
1.298 0.006
12.87 33.72
91.4
2.42.263 1.318
3.629.437 0.024
0.052 0.106
0.054 0.143
0.094 0.052
0.483 0.005
36.17 20.85
85.14
3.583.143 0.857
3.923.220 0.022
0.058 0.056
0.039 0.275
0.069 0.046
0.517 0.007
50.04 13.45
4.518.237 0.020
0.872 0.094
81.91
7.833.111 0.508
0.431 0.028
0.027 0.116
0.052 0.038
49.54 7.98
5.510.698 0.011
1.195 0.014
0.011 1.156
0.024 0.040
0.516 0.271
65.26
32.173.606 0.286
57.29 4.48
6.58.140 0.011
0.575 0.013
0.006 0.523
0.014 0.028
0.242 0.129
47.85
46.614.131 0.130
65.48 2.03
Table B2 continued.
Appendix B page 230 1
Laser
Isotope Ratios
Power
(%)
40
Ar/39Ar
38
Ar/39Ar
37
Ar/39Ar
36
Ar/39Ar
39
-10
Ca/K
%40Ar
Cl/K
%39Ar
40
Ar*/39ArK
Age±2σ
3
LW-85 (granite), hornblende, J = 0.008954±0.000014; volume ArK = 273.18 x 10 cm , integrated age = 70.84±1.91 Ma (2σ)
0.745±0.050
0.494± 0.046
0.641±0.057
1.936
0.143
104.4
0.41-7.245±6.955
-121.01±120.15
2.339.408 0.025
0.394 0.040
0.102 0.029
0.118 0.084
0.394
0.083
85.83
1.985.054 2.794
79.85 43.18
2.626.917 0.017
0.296 0.030
0.093 0.024
0.082 0.044
0.31
0.062
86.33
2.163.209 0.992
51.11 15.57
2.920.263 0.016
1.065 0.023
0.072 0.038
0.058 0.093
0.429
0.24
79.38
2.873.648 1.568
57.98 24.53
3.210.751 0.018
1.910 0.020
0.016 0.292
0.022 0.032
0.197
0.436
55.19
14.094.557 0.222
72.16 3.44
3.58.095 0.009
1.997 0.012
-0.015 1.902
0.012 0.042
0
0.456
39.79
36.124.717 0.161
74.63 2.50
3.77.993 0.018
2.022 0.019
ERR
0.013 0.064
0
0.462
41.36
15.944.392 0.261
69.59 4.06
3.97.791 0.017
2.003 0.021
0.019 0.180
0.012 0.041
0.445
0.458
37.27
14.744.553 0.176
72.10 2.73
4.18.436 0.014
2.085 0.016
0.018 1.034
0.015 0.089
0
0.477
42.69
8.644.337 0.400
68.73 6.21
4.413.390 0.009
2.280 0.014
0.031 0.549
0.032 0.065
0
0.521
59.67
3.054.450 0.611
70.49 9.49
2180.329±0.044
39
-10
3
LW-61 (hornblende-plagioclase porphyritic dyke), hornblende, J = 0.008954±0.000014; volume ArK = 273.18 x 10 cm , integrated age = 70.84±1.91 Ma (2σ)
21318.31±0.09
3.67±0.11
2.38±0.09
4.72±0.09
15.841
0.782
103.72
0.1-54.262±67.377
-1200.38±2117.39
2.3737.79 0.12
1.40 0.14
2.67 0.13
2.74 0.13
11.731
0.256
105.26
0.09-39.216 58.106
2.6594.23 0.08
1.32 0.12
2.05 0.08
2.12 0.12
10.004
0.249
100.06
0.12-0.358 69.862
-5.79 1132.29
2.9185.90 0.10
1.98 0.11
0.92 0.10
0.71 0.12
8.326
0.452
108.34
0.32-12.368 14.924
-211.77 271.14
3.224.20 0.03
3.18 0.04
0.11 0.29
0.07 0.07
2.098
0.731
80.5
3.443.790 1.410
60.21 22.04
3.59.60 0.02
3.25 0.02
0.02 0.24
0.02 0.06
0.87
0.745
46.7
32.294.811 0.331
76.09
5.12
3.78.70 0.01
3.22 0.02
0.01 0.19
0.01 0.07
0.569
0.738
42.11
42.254.778 0.306
75.58
4.74
3.910.50 0.03
3.23 0.02
0.03 0.88
0.02 0.11
0.832
0.742
53.3
16.274.406 0.787
69.81 12.24
4.332.68 0.04
5.49 0.04
0.13 0.06
0.11 0.11
1.182
1.268
89.18
2.432.808 3.310
44.80 52.16
5.528.46 0.03
5.34 0.03
0.13 0.22
0.10 0.06
1.725
1.233
92.74
2.71.613 1.466
25.87 23.35
-780.45 1446.85
Table B2 continued.
Appendix B page 231 Figure B1 ‐ Geology map of study area showing the location of Ar‐Ar samples. Map projection is UTM Zone 17, Southern Hemisphere (WGS84). Appendix B page 232 120
Plateau steps are filled, rejected steps are open
box heights are 2σ
Plateau age = 47.1 ± 3.1 Ma
(2σ, including J-error of .16%)
MSWD = 0.85, probability=0.55
Includes 100% of the 39Ar
100
Age (Ma)
80
60
40
20
0
0
20
40
60
80
100
Cumulative 39Ar Percent
Figure B2 – Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐06 (granodiorite). 120
Plateau steps are filled, rejected steps are open
box heights are 2σ
110
100
Plateau age = 48.6 ± 7.4 Ma
(2σ, including J-error of .16%)
MSWD = 1.3, probability=0.24
Includes 75.1% of the 39Ar
90
Age (Ma)
80
70
60
50
40
30
20
10
0
0
20
40
60
Cumulative
39
80
100
Ar Percent
Figure B3 ‐ Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐07 (rhyolitic volcaniclastic) Appendix B page 233 120
Plateau steps are filled, rejected steps are open
box heights are 2σ
100
Plateau age = 52.2 ± 3.4 Ma
(2σ, including J-error of .16%)
MSWD = 0.99, probability=0.44
Includes 96% of the 39Ar
Age (Ma)
80
60
40
20
0
0
20
40
60
80
100
Cumulative 39Ar Percent
Figure B4 – Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐36 (diorite). Plateau steps are filled, rejected steps are open
120
100
Plateau age = 55.0 ± 3.8 Ma
(2σ, including J-error of .16%)
MSWD = 1.7, probability=0.18
Includes 43.6% of the 39Ar
80
Age (Ma)
box heights are 2σ
60
40
20
0
0
20
40
60
80
100
Cumulative 39Ar Percent
Figure B5 ‐ Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐88 (hornblende porphyritic dyke). Appendix B page 234 Plateau steps are filled, rejected steps are open
120
box heights are 2σ
Plateau age = 72.4 ± 1.5 Ma
(2σ, including J-error of .16%)
MSWD = 1.4, probability=0.23
Includes 95.5% of the 39Ar
100
Age (Ma)
80
60
40
20
0
0
20
40
60
80
100
Cumulative 39Ar Percent
Figure B6 ‐ Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐85 (hornblende granite). 120
Plateau steps are filled, rejected steps are open
box heights are 2σ
100
Age (Ma)
80
60
Plateau age = 74.9 ± 3.3 Ma
(2σ, including J-error of .16%)
MSWD = 1.4, probability=0.18
Includes 97.3% of the 39Ar
40
20
0
0
20
40
60
Cumulative
39
80
100
Ar Percent
Figure B7 ‐ Step‐heating cumulative percent of 39Ar released vs. age plot for sample LW‐61 (hornblende‐plagioclase porphyritic dyke). Appendix B page 235 B.3 References Renne, P.R., Swisher, C.C., III, Deino, A.L., Karner, D.B., Owens, T. and DePaolo, D.J. 1998. Intercalibration of standards, absolute ages and uncertainties in 40Ar/39Ar dating. Chemical Geology, 145: 117‐152. Sandeman, H.A., Archibald, D.A., Grant, J.W., Villeneuve, M.E., Ford, F.D. 2000. Characterization of the chemical composition and (super 40) Ar‐ (super 39) Ar systematics of intralaboratory standard MAC‐83 biotite. Current Research, Geological Survey of Canada, Report: 1999‐F, pp.13‐26. Appendix B page 236 C.1 Appendix C. Lithogeochemistry Analytical Methods, Precision, and Accuracy One hundred thirty three whole‐rock analyses were performed from samples of volcanic and sub‐volcanic rocks from the Lancones basin with emphasis on the areas of VMS deposits. Samples in the vicinity of the VMS deposits were collected mostly from core in 2000 and were analyzed at Bondar‐Clegg Laboratories (now ALS Chemex) in Vancouver, Canada using inductively coupled plasma – atomic emission spectrometry (ICP‐AES) and subsequently for additional trace elements (including rare earth elements) using inductively coupled plasma – mass spectrometry (ICP‐MS) at Memorial University, St. John’s, Canada (Table C1). Regional samples collected during 2002 from outcrop and drill core were analyzed for major and trace elements by a combination of ICP‐AES and ICP‐MS at ALS‐Chemex Laboratories in Vancouver (Table C2). Analyses at Memorial University for trace elements (REE, Y, Th, Zr, Nb, Ba, Hf and Ta) were done using ICP‐MS analysis. Rocks were pulverized at Bondar Clegg Laboratories in the process described above. Dissolution involved Na2O2 sinter technique using 10 g of material. Internal standards were used to correct for matrix and drift effects and other quality control measures included pressed powder XRF analyses. Limits of detection are reported at the 3 sigma background level. The method is described in Longerich et al. (1990) and more information is available at http://www.mun.ca/earthsciences/ICPMS/Solution_ICP‐MS.php. Data for internal standards analyzed during this study are compared to given values in table C3. Normalized trace element plots for repeat analyses of internal reference materials compared to the analytical detection limit is given in figure C1, A. Average values for the reference materials compared to accepted values are shown in on a normalized trace element plot in figure C1, B. Appendix C Page 237 Analyses at Bondar‐Clegg and ALS Chemex involved coarse crushing of rock samples to better than 70% of the sample passing 6 mm and grinding in a ring mill pulverizer using a standard low chrome steel ring set. At least 85% of the pulverized passes through a 75 micron screen. For major element oxides, a prepared sample (0.200 g) is added to lithium metaborate/lithium tetraborate flux (0.90 g), mixed well and fused in a furnace at 1000°C. The resulting melt is then cooled and dissolved in 100 mL of 4% nitric acid/2% hydrochloric acid. This solution is then analyzed by ICP‐AES and the results are corrected for spectral inter‐
element interferences. For trace elements, a prepared sample (0.200 g) is added to lithium metaborate flux (0.90 g), mixed well and fused in a furnace at 1000°C. The resulting melt is then cooled and dissolved in 100 mL of 4% nitric acid. This solution is then analyzed by inductively coupled plasma ‐ mass spectrometry. For determining loss on ignition, a prepared sample (1.0 g) is placed in an oven at 1000°C for one hour, cooled, an weighed. Further details are available from ALS Chemex at http://www.alschemex.com/learnmore/learnmore‐techinfo‐
multielement‐wholerock.htm. Data for in‐house standards analyzed in this study versus accepted values are shown in table C4. Normalized trace element plots for repeat analyses of in‐house standards compared to the analytical detection limit is given in figure C1a. Average values for the in‐house standards compared to accepted values are shown in on a normalized trace element plot in figure C1b. Appendix C Page 238 Table C1. Whole rock geochemical analyses. Major oxides are from Bondar‐Clegg Laboratories. Trace elements are from Memorial University. Abbreviations: CSLF = Cerro San Lorenzo Formation; CEEF = Cerro El Ereo Formation; LBF = La Bocana Formation. D = dacite; A = andesite; B = basalt; R = rhyolite; RD = rhyolite dyke (post mineralization); bx = breccia. Appendix C Page 239 Table C1. (Cont) Appendix C Page 240 Table C1. (Cont) Appendix C Page 241 Table C1. (Cont) Appendix C Page 242 Table C1. (Cont) Appendix C Page 243 Table C1. (Cont) Appendix C Page 244 Table C1. (Cont) Appendix C Page 245 Table C2. Whole rock geochemical analyses. Major oxides are from Bondar‐Clegg Laboratories. Trace elements are from Memorial University. Appendix C Page 246 Table C2. (cont) Appendix C Page 247 Table C2. (cont) Appendix C Page 248 Table C2. (cont) Appendix C Page 249 Table C2. (cont) Appendix C Page 250 Table C2. (cont) Appendix C Page 251 Table C2. (cont) Appendix C Page 252 Table C2. (cont) Appendix C Page 253 Table C2. (cont) Appendix C Page 254 Table C2. (cont) Appendix C Page 255 Table C3. Memorial University analyses of in‐house standards run with samples from this study. Sample
MRG-1(MUN run 78-108)
BR-688 (MUN run 161-191
MRG-AVG
BRG-AVG
Limit of Detection
BLANK-25
BLANK-26
BLANK-27
BLANK-28
BR-688-27
BR-688-25
MRG-1-25
MRG-1-27
Y
11.5
17.81
9.9912991
16.617819
0.0092306
1.9430877
2.3479125
2.424286
1.9764842
17.697234
15.538405
9.7308945
10.251704
Sample
MRG-1(MUN run 78-108)
BR-688 (MUN run 161-191
MRG-AVG
BRG-AVG
Limit of Detection
BLANK-25
BLANK-26
BLANK-27
BLANK-28
BR-688-27
BR-688-25
MRG-1-25
MRG-1-27
Gd 160
3.97
2.88
3.8189632
2.9389857
0.0104266
0.009143
0.009946
0.0101534
0.0103336
3.1231084
2.754863
3.7526247
3.8853017
Zr
107
59.15
113.71705
64.413971
0.0628715
1.4203941
1.2090626
0.8357417
0.8066353
67.612528
61.215414
112.67704
114.75705
Tb
0.52
0.48
0.5027635
0.485098
0.0039823
0.0992321
0.1484815
0.2162612
0.1376261
0.5225309
0.447665
0.4878681
0.5176589
Nb
22.3
4.87
25.255068
5.6533517
0.0238228
0.0493796
0.007367
0.0299783
0.0033698
6.071878
5.2348254
25.065669
25.444467
Ba
49.4
163.33
44.090476
160.0276
0.1546975
1.8309588
1.8103243
2.0968618
1.6830069
175.38593
144.66928
42.10709
46.073862
La
Ce
Pr
8.83
4.98
7.678709
4.7409816
0.0076891
0.4939528
0.6192861
0.6481933
0.5683987
5.1200486
4.3619146
7.4229731
7.9344449
25.8
11.55
21.623305
10.870648
0.0057409
0.135662
0.1073477
0.1368853
0.1494276
11.651453
10.089843
21.054179
22.192431
Ho
Er
3
0.49
1.16
3.21
0.7
2.1
2.7189793 0.466624 1.1822944
3.268199 0.7128201 2.2701266
0.015266 0.0029849 0.0136461
0.0027314 0.0084104 0.0030661
0.0077109
0.0018642 0.0081206
0.0011799 0.0083888
3.5320501 0.7495594 2.3861954
3.0043479 0.6760809 2.1540577
2.6569843 0.4629034 1.1875442
2.7809743 0.4703446 1.1770445
Tm
0.14
0.3
0.1327873
0.3059436
0.0054858
0.000543
Dy
0.3243691
0.2875182
0.1324349
0.1331398
Nd
3.71
1.65
3.1389724
1.5547568
0.0065007
0.0931578
0.1291485
0.1210043
0.1208334
1.6719865
1.4375272
3.0390699
3.2388748
Yb
17.6
8.03
15.492122
7.6410572
0.0302344
0.0151615
0.0122866
0.0110069
8.168316 2.479277 1.0672273
7.1137983 2.2569791 0.954518
15.046007 4.185491 1.3668896
15.938237 4.3226176 1.3996091
Lu
0.79
2
0.7418058
1.9527507
0.0154391
0.1602147
0.2101635
0.1912325
0.3672724
2.0867531
1.8187484
0.7319475
0.7516642
Sm
Eu
Gd 157
4.34
1.38
2.3
0.94
4.2540543 1.3832493
2.3681281 1.0108726
0.0189688 0.0055531 0.0178039
0.0030624
Hf
0.11
0.3
0.1039626
0.318213
0.0023617
0.005113
0.0081732
0.0066691
0.0091971
0.3387535
0.2976726
0.1028635
0.1050616
3.89
1.54
4.8753173
2.1618878
0.0174635
0.5111245
0.4856024
0.4225627
0.4107917
2.2301692
2.0936064
4.8652768
4.8853579
Ta
#N/A
#N/A
#N/A
#N/A
Th
0.74
0.18
0.3638535
0.1723741
0.0154686
0.003583
0.0028297
0.0015785
0.002473
0.1580354
0.1867129
0.4407932
0.2869139
0.82
0.33
0.6921084
0.3225851
0.0038651
0.0288197
0.0224256
0.0186191
0.0213467
0.3493235
0.2958467
0.6503028
0.7339139 Appendix C Page 256 Table C4. ALS Chemex analyses of MDRU standards run with samples from this study. SAMPLE
BAS1
BAS1
BAS1
BAS1 Average (n=3)
BAS-1 (Piercey, 2001)
Al2O3
15.36
15.52
15.61
15.50
15.12
BaO
0.02
0.02
0.02
0.02
P-1
P-1
P-1
P-1 Average (n=3)
P-1 (Piercey, 2001)
14.32
14.39
14.37
14.36
14.1
0.09
0.09
0.09
0.09
0.01
SAMPLE
BAS1
BAS1
BAS1
BAS1 Average (n=3)
BAS-1 (Piercey, 2001)
P-1
P-1
P-1
P-1 Average (n=3)
P-1 (Piercey, 2001)
LOD
LOD
CaO
8.14
8.1
8.1
8.11
8.28
Cr2O3 Fe2O3
0.03 10.83
0.03 10.76
0.03 10.85
0.03 10.81
11.16
K2O
0.53
0.54
0.55
0.54
0.56
MgO
7.19
7.14
7.19
7.17
7.35
MnO
0.13
0.13
0.14
0.13
0.14
Na2O
3.28
3.3
3.36
3.31
3.28
P2O5
0.26
0.24
0.15
0.22
0.22
SiO2
52.75
52.98
53.34
53.02
53.56
SrO
0.05
0.05
0.05
0.05
3.77
3.77
3.75
3.76
3.9
2.03
2.08
2.05
2.05
2.12
1.01
1.02
1.02
1.02
1.11
0.07
0.07
0.07
0.07
0.08
3.8
3.89
3.85
3.85
3.8
0.1
0.13
0.07
0.10
0.08
70.12
70.3
70.41
70.28
70.96
0.02
0.02
0.02
0.02
TiO2
1.22
1.23
1.24
1.23
1.31
LOI TOTAL
-0.17 99.62
-0.11 99.93
-0.27 100.35
-0.18 99.97
0.78 100.96
0.35
0.35
0.35
0.35
0.38
0.59 99.79
0.33 99.83
0.38 99.98
0.43 99.87
0.72 100.56
Ag
<1
<1
<1
0.28
Ba
180.0
187.0
184.5
183.8
194
Ce
21.0
20.5
22.0
21.2
21.8
Co
51.5
44.0
43.0
46.2
42.2
Cr
260
230
180
223
226.6
Cs
0.1
0.1
0.1
0.1
0.12
Cu
275
155
45
158
59.6
Dy
2.9
2.8
3.1
2.9
3.26
Er
1.5
1.4
1.6
1.5
1.54
Eu
1.2
1.1
1.1
1.1
1.28
Ga
20
20
18
19
19.6
<1
<1
822
798
25
25
7.5
8.5
160
180
1.2
1.3
15
15
2.9
3
1.9
2.1
0.7
0.8
15
15
0.3
810
724
25
28
8
6.2
170
149.2
1.25
1.22
15
15.5
2.95
3.34
2
2.1
0.75
0.78
15
15
3.51
3.38
3.54
3.48
3.49
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
0.01
1
0.5
0.5
0.5
10
0.1
5
0.1
0.1
0.1
1
Gd
3.4
3.3
3.4
3.4
3.8
Hf
2
2
2
2
2.36
Ho
0.6
0.6
0.6
0.6
0.61
La
8.5
8.5
11.0
9.3
9.28
Lu
0.2
0.1
0.1
0.1
0.21
Mo
8
8
6
7
2.24
Nb
8
7
7
7
8.2
Nd
13.0
13.5
13.0
13.2
13.6
Ni
190
175
160
175
172
Pb
20
5
5
10
4.4
Pr
2.9
2.9
2.8
2.9
3
Rb
7.6
7
7.4
7.3
6.96
Sm
3.4
3.4
3.1
3.3
3.5
Sn
1
1
1
1
1.2
Sr
483
490
404
459
502
Ta
0.5
<0.5
0.5
0.5
0.45
Tb
0.5
0.5
0.5
0.5
0.57
Th
1
1
<1
1
0.84
Tl
<0.5
<0.5
<0.5
U
<0.5
<0.5
<0.5
0.32
V
200
190
165
185
152
W
2
2
2
2
0.06
Tm
0.2
0.2
0.2
0.2
0.23
Y
15.5
15
16
15.5
18.4
Yb
1.4
1.3
1.2
1.3
1.4
Zn
120
105
110
112
91.4
Zr
84.5
80
79.5
81
94.5
2.7
3.1
4
4
0.7
0.7
11.5
11.5
0.4
0.4
6
6
4
5
12.5
12.5
5
10
5
15
3.1
3.2
51.4
52.4
2.9
2.9
3
2
247
244
<0.5
0.5
0.5
0.5
5
5
<0.5
<0.5
0.4
0.3
1.5
2
70
85
4
2
19.5
20.5
2.3
2.3
50
50
131.5
142.5
2.9
3.12
4
3.76
0.7
0.72
11.5
13.2
0.4
0.4
6
0.52
4.5
3.78
12.5
13
7.5
10
10.2
3.15
3.36
51.9
50.4
2.9
2.92
2.5
2.44
245.5
256
0.5
0.3
0.5
0.52
5
4.38
0.35
0.35
1.75
1.48
77.5
58.2
3
0.31
20
22.8
2.3
2.46
50
44
137
126
0.1
1
0.1
0.5
0.1
2
1
0.5
5
5
0.1
0.2
0.1
1
0.1
0.5
0.1
1
0.5
0.1
0.5
5
1
0.5
0.1
5
0.5
Appendix C Page 257 Figure C8. (following page). Primitive mantle‐normalized trace element diagrams for repeat analyses of in‐house and internal reference material conducted during this study. A. ICP‐MS data (Memorial Univ.) for standards MRG‐1 and BR‐688 compared to the detection limit. B. Average analyses for the reference materials from this study in (A) compared to given values from previous analysis by Memorial Univ. of the material. C. ICP‐AES data (ALS Chemex) for MDRU reference samples BAS‐1 and P‐1 as compared to the detection limit for the technique. D. Average values for the reference materials from this study from in (C) compared to data from Piercey (2001). Appendix C Page 258 Appendix C Page 259 C.2 References Longerich, H.P., Jenner, G.A., Fryer, B.J. and Jackson, S.E., 1990. Inductively coupled plasma‐
mass spectrometric analysis of geological samples: an overview. Chemical Geology 83: 105‐
118. Piercey, S.J. 2001. Petrology and tectonic setting of felsic and mafic volcanic and intrusive rocks in the Finlayson Lake volcanic‐hosted massive sulphide (VHMS) district, Yukon Canada: A record of mid‐Paleozoic arc and back‐arc magmatism and metallogeny. Unpublished Ph.D. thesis, The University of British Columbia, 304 p. Appendix C Page 260