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This chapter discusses how radiogenic and stable isotopes can be used in the study of metallic mineral deposits. Although the chapter is mostly focused on the radiogenic (Pb, Os) and heavy stable (Fe, Cu, Zn) isotopes of metallic elements, we complement the discussion highlighting also the power of stable isotopes of light elements, which are major to significant components of hydrothermal fluids and rocks (e.g., H, B, C, N, O, S), as well as of radiogenic isotopes of elements (Sr, Nd, Hf ) that are useful in tracing fluid/magma sources and their interaction with the host rocks. In the first part of this chapter we discuss general aspects of isotopes clarifying the differences between stable non-radiogenic and stable radiogenic isotopes and, consequently, their different applicability to metallogenic studies. Due to their properties, stable non-radiogenic isotopes record mass-dependent fractionation that occur in many reactions associated with the formation of mineral deposits. Mass-dependent fractionation of stable non-radiogenic isotopes occurs both under equilibrium and non-equilibrium (kinetic) conditions of the reactions leading to ore mineral deposition and is controlled by various physico-chemical parameters, like, among the principal ones, temperature, oxygen fugacity, and biological activity. Therefore, stable non-radiogenic isotopes can inform us about the physico-chemical and, eventually, biological processes that control ore mineral deposition and also on the sources of some metals (e.g., transition metal isotopes of elements like Fe, Cu, Zn) or of the fluids (e.g., H, C, O, N, S isotopes) and even of metal ligands (e.g., S, Cl). We conclude the first part of the chapter providing some hints on the strategy of sampling and on the instrumentation related to isotopic studies. In the second part we discuss radioactive-radiogenic isotope systems and their applications in metallogenic studies of metallic mineral deposits. Stable radiogenic isotopes are characterized by relative variations that are controlled, in each geological system, by the addition of a radiogenic component of an isotope, derived from the decay of a radioactive parent, to the same radiogenic isotope already present in the Earth since its formation  4.55 Gyr ago. This relative variation is usually expressed as the ratio of a radiogenic isotope of an element to a non-radiogenic isotope of the same element. The ratio of these two isotopes has increased since the Earth formation and the magnitude of its variations depends on the radioactive/ radiogenic isotope ratios in different geological systems and on the time elapsed since the system has formed. The Earth is  4.55 Gyr old and has evolved from an initially homogeneous isotopic composition to reservoirs (e.g., mantle, crust) and crustal rocks with very variable radioactive/radiogenic isotope ratios due to magmatic, metamorphic, weathering, atmospheric and biologic processes, among others. This has resulted in extremely large variations of radiogenic isotopes in rocks and reservoirs of the Earth which can track various geological processes. In ore geology, stable radiogenic isotopes are best suited for tracing metal (e.g., Pb, Os) sources from different rocks and reservoirs (e.g., mantle, upper crust, lower crust), fluid-rock interactions (i.e., the hydrothermal plumbing system), or magma-host rock interactions (e.g., host rock assimilation by magmas associated with magmatic-hydrothermal deposits). Radioactive-radiogenic isotope systems allow us to determine also absolute ages of suitable minerals that are found in mineral deposits. This is an essential information in metallogeny that allows us to link the formation of a mineral deposit to a specific geological process and/or to specific periods of the Earth’s history. We discuss various dating methods that are extensively applied to date mineral deposits. These methods can be subdivided into those that allow a direct dating of ore minerals (e.g., RedOs dating of molybdenite, UdPb dating of cassiterite) and those that allow dating of minerals that are demonstrably related with the mineralization (e.g., UdPb dating of zircon from magmatic rocks associated with magmatic-hydrothermal deposits; Ar/Ar dating of K-bearing minerals resulting from alteration associated with various types of mineral deposits). We discuss pros and cons of using these various methods and also mention methods that are less used (because potentially less accurate and precise), but sometimes represent the only possibility to provide an age to deposit types that are notoriously difficult to date (e.g., MVT and Carlin-type deposits). We highlight the power of both stable radiogenic and non-radiogenic isotopes in unravelling the genesis of metallic mineral deposits through a series of conceptual and real examples applied to a broad range of mineral deposit types such as porphyry systems (i.e., porphyry deposits, high- and intermediate-sulfidation epithermal deposits, skarn, carbonate replacement deposits, sediment-hosted Au deposits), low-sulfidation epithermal deposits, IOCG deposits, ortho-magmatic deposits, volcanic-hosted massive sulfide deposits (VHMS), sediment-hosted deposits (stratiform copper, MVT), and supergene deposits. In the third part of the chapter, we discuss the use of transition metal stable non-radiogenic isotopes to mineral deposits. Although in its infancy, the application of transition metal isotopes to mineral deposit investigation is quickly growing because these isotopes allow us to address different aspects of the formation of mineral deposits compared to radiogenic isotopes. In particular, isotopes of transition metals (like stable isotopes of light elements) undergo mass-dependent fractionation processes that may be associated with different types of equilibrium and non-equilibrium chemical, physical and biological reactions occurring during the formation of mineral deposits. We focus on the applications of the isotopes of Cu, Fe and Zn to various deposit types, because isotopes of these transition metals are those that have been most extensively used in mineral deposit studies. Mass-independent fractionation may also occur for isotopes of some elements and could be a developing field that has not yet been extensively explored in the study of mineral deposits. 

Iron oxide-apatite (IOA) deposits, also known as magnetite-apatite or Kiruna-type deposits, are a major source of iron and potentially of rare earth elements and phosphorus. To date, the youngest representative of this group is the Pleistocene (~2 Ma) El Laco deposit, located in the Andean Cordillera of northern Chile. El Laco is considered a unique type of IOA deposit because of its young age and its volcanic-like features. Here we report the occurrence of similarly young IOA-type mineralization hosted within the Laguna del Maule Volcanic Complex, an unusually large and recent silicic volcanic system in the south-central Andes. We combined field observations and aerial drone images with detailed petrographic observations, electron microprobe analysis (EMPA), and 40Ar/39Ar dating to characterize the magnetite mineralization—named here “Vetas del Maule”—hosted within andesites of the now extinct La Zorra volcano (40Ar/39Ar plateau age of 1.013 ± 0.028 Ma). Five different styles of magnetite mineralization were identified: (1) massive magnetite, (2) pyroxene-actinolite-magnetite veins, (3) magnetite hydrothermal breccias, (4) disseminated magnetite, and (5) pyroxene-actinolite veins with minor magnetite. Field observations and aerial drone imaging, coupled with microtextural and microanalytical data, suggest a predominantly hydrothermal origin for the different types of mineralization. 40Ar/39Ar incremental heating of phlogopite associated with the magnetite mineralization yielded a plateau age of 873.6 ± 30.3 ka, confirming that the emplacement of Vetas del Maule postdated that of the host andesite rocks. Our data support the hypothesis that the magnetite mineralization formed in a volcanic setting from Fe-rich fluids exsolved from a magma at depth. Ultimately, Vetas del Maule provides evidence that volcanic-related IOA mineralization may be more common than previously thought, opening new opportunities of research and exploration for this ore deposit type in active volcanic arcs. 

The Coastal Ocean Environment Summer School In Nigeria and Ghana (COESSING; https://coessing.org) has been run for one week every year since 2015. The school, an endorsed project of the United Nations Decade of Ocean Science for Sustainable Development (2021–2030), has provided a platform for approximately 1,000 scientists from Africa, the United States, and Europe to exchange scientific knowledge, to network, to learn, and to collaborate. Our interdisciplinary, multicultural, and multi-institutional approach offers a model for knowledge exchange across the globe and across different educational levels. 

Abstract The mineral apatite, Ca10(PO4)6(F,OH,Cl)2, incorporates sulfur (S) during crystallization from S-bearing hydrothermal fluids and silicate melts. Our previous studies of natural and experimental apatite demonstrate that the oxidation state of S in apatite varies systematically as a function of oxygen fugacity (fO2). The S oxidation states –1 and –2 were quantitatively identified in apatite crystallized from reduced, S-bearing hydrothermal fluids and silicate melts by using sulfur K-edge X-ray absorption near-edge structure spectroscopy (S-XANES) where S 6+/ΣS in apatite increases from ~0 at FMQ-1 to ~1 at FMQ+2, where FMQ refers to the fayalite-magnetite-quartz fO2 buffer. In this study, we employ quantum-mechanical calculations to investigate the atomistic structure and energetics of S(-I) and S(-II) incorporated into apatite and elucidate incorporation mechanisms. One S(-I) species (disulfide, S22−) and two S(-II) species (bisulfide, HS−, and sulfide, S2−) are investigated as possible forms of reduced S species in apatite. In configuration models for the simulation, these reduced S species are positioned along the c-axis channel, originally occupied by the column anions F, Cl, and OH in the end-member apatites. In the lowest-energy configurations of S-incorporated apatite, disulfide prefers to be positioned halfway between the mirror planes at z = 1/4 and 3/4. In contrast, the energy-optimized bisulfide is located slightly away from the mirror planes by ~0.04 fractional units in the c direction. The energetic stability of these reduced S species as a function of position along the c-axis can be explained by the geometric and electrostatic constraints of the Ca and O planes that constitute the c-axis channel. The thermodynamics of incorporation of disulfide and bisulfide into apatite is evaluated by using solid-state reaction equations where the apatite host and a solid S-bearing source phase (pyrite and Na2S2(s) for disulfide; troilite and Na2S(s) for sulfide) are the reactants, and the S-incorporated apatite and an anion sink phase are the products. The Gibbs free energy (ΔG) is lower for incorporation with Na-bearing phases than with Fe-bearing phases, which is attributed to the higher energetic stability of the iron sulfide minerals as a source phase for S than the sodium sulfide phases. The thermodynamics of incorporation of reduced S is also evaluated by using reaction equations involving dissolved disulfide and sulfide species [HnS(aq)(2−n) and HnS(aq)(2−n); n = 0, 1, and 2] as a source phase. The ΔG of S-incorporation increases for fluorapatite and chlorapatite, and decreases for hydroxylapatite, as these species are protonated (i.e., as n changes from 0 to 2). These thermodynamic results demonstrate that the presence of reduced S in apatite is primarily controlled by the chemistry of magmatic and hydrothermal systems where apatite forms (e.g., an abundance of Fe; solution pH). Ultimately, our methodology developed for evaluating the thermodynamics of S incorporation in apatite as a function of temperature, pH, and composition is highly applicable to predicting the trace and volatile element incorporation in minerals in a variety of geological systems. In addition to solid-solid and solid-liquid equilibria treated here at different temperatures and pH, the methodology can be easily extended to different pressure conditions by just performing the quantum-mechanical calculations at elevated pressures. 

Ti-isotope fractionation on the most Ti-rich minerals on Earth has not been reported. Therefore, we present a chemical preparation and separation technique for Ti-rich minerals for mineralogic, petrologic, and economic geologic studies. A two-stage ion-exchange column procedure modified from the previous literature is used in the current study to separate Ti from Fe-rich samples, while α-TiO2 does not require chemical separation. Purified solutions in conjunction with solution standards were measured on two different instruments with dry plasma and medium-resolution mode providing mass-dependent results with the lowest errors. 49/47TiOL-Ti for the solution and solids analyzed here demonstrate a range of >5‰ far greater than the whole procedural 1 error of 0.10‰ for a synthetic compound and 0.07‰ for the mineral magnetite; thus, the procedure produces results is resolvable within the current range of measured Ti-isotope fractionation in these minerals. 

Abstract Magnetite is the main constituent of iron oxide–apatite (IOA) deposits, which are a globally important source of Fe and other elements such as P and REE, critical for modern technologies. Geochemical studies of magnetite from IOA deposits have provided key insights into the ore-forming processes and source of mineralizing fluids. However, to date, only qualitative estimations have been obtained for one of the key controlling physico-chemical parameters, i.e., the temperature of magnetite formation. Here we reconstruct the thermal evolution of Andean IOA deposits by using magnetite thermometry. Our study comprised a > 3000 point geochemical dataset of magnetite from several IOA deposits within the Early Cretaceous Chilean Iron Belt, as well as from the Pliocene El Laco IOA deposit in the Chilean Altiplano. Thermometry data reveal that the deposits formed under a wide range of temperatures, from purely magmatic (~ 1000 to 800 °C), to late magmatic or magmatic-hydrothermal (~ 800 to 600 °C), to purely hydrothermal (< 600 °C) conditions. Magnetite cooling trends are consistent with genetic models invoking a combined igneous and magmatic-hydrothermal origin that involve Fe-rich fluids sourced from intermediate silicate magmas. The data demonstrate the potential of magnetite thermometry to better constrain the thermal evolution of IOA systems worldwide, and help refine the geological models used to find new resources. 

Abstract Iron oxide copper-gold (IOCG) and iron oxide-apatite (IOA) deposits are major sources of Fe, Cu, and Au. Magnetite is the modally dominant and commodity mineral in IOA deposits, whereas magnetite and hematite are predominant in IOCG deposits, with copper sulfides being the primary commodity minerals. It is generally accepted that IOCG deposits formed by hydrothermal processes, but there is a lack of consensus for the source of the ore fluid(s). There are multiple competing hypotheses for the formation of IOA deposits, with models that range from purely magmatic to purely hydrothermal. In the Chilean iron belt, the spatial and temporal association of IOCG and IOA deposits has led to the hypothesis that IOA and IOCG deposits are genetically connected, where S-Cu-Au–poor magnetite-dominated IOA deposits represent the stratigraphically deeper levels of S-Cu-Au–rich magnetite- and hematite-dominated IOCG deposits. Here we report minor element and Fe and O stable isotope abundances for magnetite and H stable isotope abundances for actinolite from the Candelaria IOCG deposit and Quince IOA prospect in the Chilean iron belt. Backscattered electron imaging reveals textures of igneous and magmatic-hydrothermal affinities and the exsolution of Mn-rich ilmenite from magnetite in Quince and deep levels of Candelaria (>500 m below the bottom of the open pit). Trace element concentrations in magnetite systematically increase with depth in both deposits and decrease from core to rim within magnetite grains in shallow samples from Candelaria. These results are consistent with a cooling trend for magnetite growth from deep to shallow levels in both systems. Iron isotope compositions of magnetite range from δ56Fe values of 0.11 ± 0.07 to 0.16 ± 0.05‰ for Quince and between 0.16 ± 0.03 and 0.42 ± 0.04‰ for Candelaria. Oxygen isotope compositions of magnetite range from δ18O values of 2.65 ± 0.07 to 3.33 ± 0.07‰ for Quince and between 1.16 ± 0.07 and 7.80 ± 0.07‰ for Candelaria. For cogenetic actinolite, δD values range from –41.7 ± 2.10 to –39.0 ± 2.10‰ for Quince and from –93.9 ± 2.10 to –54.0 ± 2.10‰ for Candelaria, and δ18O values range between 5.89 ± 0.23 and 6.02 ± 0.23‰ for Quince and between 7.50 ± 0.23 and 7.69 ± 0.23‰ for Candelaria. The paired Fe and O isotope compositions of magnetite and the H isotope signature of actinolite fingerprint a magmatic source reservoir for ore fluids at Candelaria and Quince. Temperature estimates from O isotope thermometry and Fe# of actinolite (Fe# = [molar Fe]/([molar Fe] + [molar Mg])) are consistent with high-temperature mineralization (600°–860°C). The reintegrated composition of primary Ti-rich magnetite is consistent with igneous magnetite and supports magmatic conditions for the formation of magnetite in the Quince prospect and the deep portion of the Candelaria deposit. The trace element variations and zonation in magnetite from shallower levels of Candelaria are consistent with magnetite growth from a cooling magmatic-hydrothermal fluid. The combined chemical and textural data are consistent with a combined igneous and magmatic-hydrothermal origin for Quince and Candelaria, where the deeper portion of Candelaria corresponds to a transitional phase between the shallower IOCG deposit and a deeper IOA system analogous to the Quince IOA prospect, providing evidence for a continuum between both deposit types.