Artur Deditius

Iron oxide copper-gold (IOCG) and Kiruna-type iron oxide-apatite (IOA) deposits are commonly spatially and temporally associated with one another, and with coeval magmatism. Here, we use trace element concentrations in magnetite and pyrite, Fe and O stable isotope abundances of magnetite and hematite, H isotopes of magnetite and actinolite, and Re-Os systematics of magnetite from the Los Colorados Kiruna-type IOA deposit in the Chilean iron belt to develop a new genetic model that explains IOCG and IOA deposits as a continuum produced by a combination of igneous and magmatic-hydrothermal processes. The concentrations of [Al + Mn] and [Ti + V] are highest in magnetite cores and decrease systematically from core to rim, consistent with growth of magnetite cores from a silicate melt, and rims from a cooling magmatic-hydrothermal fluid. Almost all bulk δ 18 O values in magnetite are within the range of 0 to 5‰, and bulk δ 56 Fe for magnetite are within the range 0 to 0.8‰ of Fe isotopes, both of which indicate a magmatic source for O and Fe. The values of δ 18 O and δD for actinolite, which is paragenetically equivalent to magnetite, are, respectively, 6.46 ± 0.56 and-59.3 ± 1.7‰, indicative of a mantle source. Pyrite grains consistently yield Co/Ni ratios that exceed unity, and imply precipitation of pyrite from an ore fluid evolved from an intermediate to mafic magma. The calculated initial 187 Os/ 188 Os ratio (Osi) for magnetite from Los Colorados is 1.2, overlapping Osi values for Chilean porphyry-Cu deposits, and consistent with an origin from juvenile magma. Together, the data are consistent with a geologic model wherein (1) magnetite microlites crystallize as a near-liquidus phase from an intermediate to mafic silicate melt; (2) magnetite microlites serve as nucleation sites for fluid bubbles and promote volatile saturation of the melt; (3) the volatile phase coalesces and encapsulates magnetite microlites to form a magnetite-fluid suspension; (4) the suspension scavenges Fe, Cu, Au, S, Cl, P, and rare earth elements (REE) from the melt; (5) the suspension ascends from the host magma during regional extension; (6) as the suspension ascends, originally igneous mag-netite microlites grow larger by sourcing Fe from the cooling magmatic-hydrothermal fluid; (7) in deep-seated crustal faults, magnetite crystals are deposited to form a Kiruna-type IOA deposit due to decompression of the magnetite-fluid suspension; and (8) the further ascending fluid transports Fe, Cu, Au, and S to shallower levels or lateral distal zones of the system where hematite, magnetite, and sulfides precipitate to form IOCG deposits. The model explains the globally observed temporal and spatial relationship between magmatism and IOA and IOCG deposits, and provides a valuable conceptual framework to define exploration strategies.

Iron oxide-apatite (IOA) and iron oxide-copper-gold deposits (IOCG) are often spatially and temporally related with one another and with coeval magmatism. However, a genetic model that accounts for observations of natural systems remains elusive, with few observational data able to distinguish among working hypotheses that invoke meteoric fluid, magmatic-hydrothermal fluid, and immiscible melts. Here, we use high-resolution trace element concentrations in magnetite, hematite and pyrite, high-precision Fe and O stable isotope data of magnetite and hematite grains, δD of magnetite and actinolite, and Re and Os in magnetite and pyrite from the Los Colorados IOA and Candelaria and Mantoverde IOCG deposits in the Chilean Iron Belt to elucidate the origin of IOA and IOCG systems. At Los Colorados, Ti, V, Al, and Mn are enriched in magnetite cores and decrease systematically from core to rim, a trend consistent with magmatic and/or magmatic-hydrothermal magnetites. High Co/Ni ratios of pyrite from Los Colorados are also consistent with a magmatic-hydrothermal origin. δD values for magnetite and actinolite indicate a mantle source for H. Values of d56Fe and d18O for magnetite and hematite from all deposits indicate a magmatic source for Fe and O. The Re-Os systematics overlap data from Andean porphyry Cu-Mo deposits and are consistent with a magmatic-hydrothermal origin. Together, the data are consistent with a genetic model wherein 1) magnetite cores crystallize from silicate melt; 2) these magnetite crystals are nucleation sites for aqueous fluid that exsolves and scavenges Fe, P, S, Cu, Au from silicate melt; 3) the magnetite-fluid suspension is less dense that the surrounding magma, allowing ascent; 4) as the suspension ascends, magnetite grows in equilibrium with the fluid and takes on a magmatic-hydrothermal character (i.e., lower Al, Mn, Ti, V); 5) during ascent, magnetite, apatite and actinolite are deposited to form IOA deposits; 6) the further ascending fluid transports Fe, Cu, Au and S toward the surface where hematite, magnetite and sulfides precipitate to form IOCG deposits. This model is globally applicable and explains the observed temporal and spatial relationship between magmatism and formation of IOA and IOCG deposits.

Arsenic-bearing pyrite, known as arsenian pyrite, is the most important and in many cases the only Au-bearing mineral in Carlin-type and some epithermal deposits. Because it incorporates Au from solutions that are undersaturated with respect to native Au, precipitation of arsenian pyrite is the key to removal of Au from solution and formation of an economic deposit. Recent micro-scale studies of arsenian pyrite from a wide range of deposits have provided new information on its role as a host for Au and on the nature and significance of its growth zoning. Two types of arsenian pyrite have been recognized so far: 1) As (super 1-) -pyrite, in which As (super 1-) substitutes for S; it is found in Carlin-type and low-sulfidation epithermal deposits that formed in reduced hydrothermal systems. 2) As (super 3+) -pyrite, in which As (super 3+) and possibly other forms of oxidized As substitute for Fe; it is found in high-sulfidation epithermal deposits that formed in oxidized hydrothermal systems. Some arsenian pyrite also contains As (super 0) in nanoscale inclusions of amorphous (originally liquid) As-Fe-S. Although this As (super 0) is not part of the pyrite structure, it does contribute to the total As content of the pyrite. Most Au in arsenian pyrite is invisible, even in high-resolution TEM observations, and it is thought to be present in the crystal lattice of the pyrite. Entry of Au into the pyrite lattice is facilitated by As. Analytical compilations for Carlin-type deposits show that As and Au concentrations in As (super 1-) -pyrite plot in a wedge-shaped zone in Au-As space with an upper concentration (C) of Au defined by: C (sub Au) =0.02.C (sub As) +4X10 (super -5) , indicating a maximum Au/As molar ratio of approximately 0.02. Pyrite with As:Au ratios above 0.02 contains discrete grains of Au, usually in nanoscale particles. At least some As (super 3+) -pyrite appears to contain Au in solid solution above this solubility limit. Other elements that are found in arsenian pyrite in relatively high (ppm to low percent) concentrations include Ag,Bi, Te, Sb, Hg and Pb; of these, at least Sb appears to show solubility relations similar to those of Au. Trace elements are not homogeneously distributed in pyrite and, instead, form both sectoral and growth zones that show a wide range of concentrations. The growth zones are concentric and occupy all faces of a crystal; they can be correlated to provide an indication of hydrothermal flow patterns and connectivity, as well as the number and composition of fluid pulses that fed the hydrothermal systems. Arsenian pyrite varies in grain size from nanoscale to large crystals; the combined presence of As and smaller grain sizes cause the pyrite to decompose more rapidly, a factor that should aid in metallurgical processing. The ability of arsenian pyrite to scavenge Au from undersaturated solutions makes it the key factor in formation of many Au deposits. In exploration for Au deposits, attention should be given to factors that might cause deposition of arsenian pyrite.

Pyrite, FeS2, the most common sulfide mineral in ore deposits, exhibits trace-element rich growth zoning that reflects chemical and/or textural changes during its formation (e.g., Large et al., 2008; Barker et al., 2009; Deditius et al., 2009a). TEM data, coupled with EMPA analyses, allow identification and characterization of structural and chemical discontinuities between growth zones in pyrite at the nanoscale. EMPA analyses and SEM observations of pyrite from high-sulfidation epithermal deposits (Yanacocha and Pueblo Viejo) reveal that highest concentrations of trace elements (Au, Ag, As, Pb, Cu, Sb, Ni, Te) coincide with porous growth zones that vary in thickness from ̴50 nm to hundreds of micrometers. HRTEM and HAADF-STEM observations show that the growth zones consist of nanolayers containing homo- or heterogeneously distributed As and Cu in single crystals of pyrite or randomly distributed (As,Au,Pb)-rich aggregates of nanoparticulate pyrite, with individual grains ranging from 8 to 900 nm in size. �False� growth zoning is formed by densely distributed crystalline or amorphous non-pyrite phases.

As-bearing pyrite is one of the main hosts for Au and other trace elements in epithermal, Carlin and mesothermal (orogenic) Au deposits. A review of our own and published SIMS, EMPA, LA-ICP-MS and PIXE analyses of pyrite from these deposits suggests that the solubility of Ag, Te, Hg, Sb and Pb in arsenian pyrite is controlled by As-content in a manner similar to that previously reported for Au by Reich et al., (2005). The trace elements can be divided into two groups that exhibit different solubility limits: i) Au, Ag, Te, Hg and Bi ii) Sb and Pb. HRTEM and HAADF-STEM observations reveal nanoparticles with compositions of Sb-As-Fe-Ni, Sb-Pb-Te, Pb-Bi, PbS and Ag in arsenian pyrite above the solubility limit. Most nanoparticles are between 5 and 200 nm, with some containing Pb reaching 500 nm. Pyrite from Carlin-type and epithermal deposits contains larger amounts of Sb and/or As than pyrite from higher-temperature orogenic gold/mesothermal deposits. This suggests that the solubility of trace elements in pyrite appears to decrease with increasing temperature.

We document the occurrence of inclusions of a (Ag,I)-rich mineral in supergene chalcocite from the Mantos de la Luna argentiferous stratabound Cu depos it in the Coastal Range of northern Chile. In this deposit, located 30 km south of Tocopilla, Cu mineralization occurs preferentially in the lower levels of amygdaloidal and porphyritic horizons. Mineral paragenesis is simple and composed exclusively of Ag-bearing supergene chalcocite (digenite), atacamite, and chrysocolla. EMPA observations reveal the presence of discrete, micron-sized (1-10 μm) inclusions of a Ag iodide mineral in supergene chalcocite. The inclusions were identified as iodargyrite by means of EDS and WDS elemental mapping. The Ag concentrations in the inclusions vary from 1.0-67.6 wt% and they are contaminated by Cu and S from chalcocite. The small size and the beam-sensitivity of the Ag-I inclusions precluded the precise descript ion of its chemical formula. However, the Ag and I elementa l maps strongly correlate with the inclusions, whereas the WD S maps of Cu and S correlate well with the chalcocite sulfide host. The occurrence of iodargyrite inclusions in supergene chalcocite suggests the involvement of iodine-rich waters during supergene enrichment at the Mantos de la Luna Cu deposit. Considering the fact that the occurrence of iodargyrite is restricted to extremely arid environments [1], our observations strongly suggest the prevalence of hyperarid conditions during the latest stag es of supergene enrichment of the Mantos de la Luna argentiferous Cu deposit in northern Chile. This suggests that supergene enrichment processes of Cu deposits in the hyperarid Atacama Desert are dynamic in nature and do not exclusively require the presence of meteoric water. Further studies are needed not only to address the isotopic signature (and age) of iodine-rich waters involved in supergene enrichment of these deposits (e.g. deep formation waters), but also to constrain the origin of iodine in the extensive nitrate deposits occurr ing in the eastern flank of the Coastal Range.

P-rich coffinite, U(Si,P)04•H20 , from the natural nuclear reactors in Bangombe, Gabon, is an important phase that incorporates percent levels of actinides and fission products. We have examined sample BAX03 (depth 12.2-12.3m) from Bangombe in order to understand micro- and nano-scale crystalo-chemical properties of P-coffinite. Electron microprobe analysis (EMPA) was completed only on coffinite inclusions (-100 fliil in size) in quartz to minimize the effect of alteration to U(VI)-phosphates and - sulfates. Based on the Si/P ratios three different chemical compositions; i) coffinite (<1.45 wt.% of P20 5) without uraninite inclusions, ii) P-coffinite, and iii) Si-ningyoite, (U,Ca,Ceh(P04)2•1-2H20. Phases ii) and iii) have inclusions of uraninite. The composition of coffinite i) is expressed to be CUo.79Cao.osREE+ Yo.04)o. s7(Si,.03Po.os)J.0704. The amount of (Y+REE)203 is <1.9 wt.%. Phosphorous substitutes for Si as evidenced by its negative correlation with Si (R2=0.87) and positive correlation with Y +REE (R2=0.6). The formula of P-coffinite (ii) is (Uo.7J-o. ss Cao.o6-o.uREE+ Yo.o7-o.Js)o.92- J.os(Sio. 39-o.s9Po.2s-o.4So.03-0.J2)o.s-o.9604, and the P20 s and (REE+Y)203 m·e as high as 9.3 wt.% and 8.64 wt.%, respectively. There is a positive correlation between Ca and P (R2=0.62), but no correlation among P, Y+REE and Si. The chemical formula of Si-rich ningyoite is (U~.43_u 3REE+ Y0.2_ o.3C

Pyrite from the Pueblo Viejo high-sulfidation deposit provides evidence for decoupling of Cu and As in hydrothermal solutions. The pyrite that shows this decoupling is in the late-stage veins that also contain sphalerite and minor enargite. Pyrite in the veins shows growth zoning that varies in composition with depth into the deposit. Deepest veins (>150m below the present surface) contain fine-grained (<211m) pyrite with 0.4% As, 0.5% Pb, 0.2% Cu, 0.1% Ag, 0.09% Te and 0.06% Sb (wt% by EMPA). At depths of 120- 103m below the present surface, pyrite contains alternating growth zones with either Cu (<0.78%) or As (<0.69%), but never both. Farther upward in the deposit concentrations of Cu and As in the two types of pyrite increase and Pb (<1.8%), Sb (<0.33%), Ag (<0.1%) and Te (<0.08%) are also in As-rich zones. At a depth of - 20m, Cu and As reach concentrations of up to 3 wt% in separate, alternating growth zones. EMPA elemental maps of the shallowest pyrites reveal that increased concentrations of As and Cu coincide spatially with decreasing concentrations of Fe and show no relation to S, suggesting that both elements substitute for Fe. Chemical compositions of Cu-pyrite and As-pyrite are: (Fe0.95CUo 06) 101S2 and (Fe09~s0_05) 101 Sb respectively. HRTEM observations on pyrite with highest Cu and As concentrations reveal that the pyrite consists of single crystals that are continuous from Cu-rich to As-rich growth zones. There is no visible (by TEM) grain boundary between Cu-rich and As-rich zones. Cu-rich growth zones contain no Cubearing inclusions, whereas As-rich growth zones contain numerous ordered nano-domains rich in As. The alternating sequence of Cu-rich and As-rich zones appears to reflect separation of Cu from As during evolution of the hydrothermal fluids. Similar decoupling of As and Cu is seen in analyses of fumaroles and fluid inclusions (both vapor and liquid), which are enriched in As and Cu, respectively. This suggests that the decoupling is related to magmatic processeses, probably involving vapor-liquid transitions.

Uranium-(VI) phases are the primary alteration products of the UO 2 in spent nuclear fuel and the UO 2+x in natural uranium deposits. The U(VI)-phases generally form sheet structures of edge-sharing UO 2 2+ polyhedra. The complexity of these structures offers numerous possibilities for coupled-substitutions of trace metals and radionuclides. The incorporation of radionuclides into U(VI)-structures provides a potential barrier to their release and transport in a geologic repository that experiences oxidizing conditions. In this study, we have used natural samples of UO 2+x , to study the U(VI)-phases that form during alteration and to determine the fate of the associated trace elements.