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1. Applications of radiogenic and transition metal isotopes to the study of metallic mineral deposits

   https://doi.org/10.1016/b978-0-323-99762-1.00010-3  

   Chiaradia, Massimo; Mathur, Ryan; Vennemann, Torsten; Simon, Adam (January 2025, Elsevier)

   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. 

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2. Redox and mineral controls on Fe and Ti isotopic fractionations during calc-alkaline magmatic differentiation

   https://doi.org/10.1016/j.gca.2023.06.016  

   Johnson, Aleisha C.; Zhang, Zhe J.; Dauphas, Nicolas; Rudnick, Roberta L.; Foden, John D.; Toc, Magali (August 2023, Geochimica et Cosmochimica Acta)

   Titanium and Fe isotopic compositions of lavas from a calc-alkaline differentiation suite and corresponding mineral separates from the Rindjani Volcano, Indonesia show that Fe and Ti isotopic fractionations between minerals and melts are lower than those recorded in other suites at all stages of differentiation. The limited isotopic fractionation for Ti is likely due to low-Ti magnetite and clinopyroxene being the dominant carriers of Ti in Rindjani lavas, as these minerals are thought to have limited equilibrium Ti isotopic fractionation relative to silicate magmas. Other magmatic differentiation suites controlled by removal of Ti-rich magnetite and characterized by a lesser role of clinopyroxene have larger Ti isotopic fractionations. This effect is an indirect consequence of the elevated Fe3+/Fe2+ ratio of calc-alkaline magmas such as Rindjani, which promotes Fe3+ incorporation into magnetite at the expense of Fe2+-Ti4+ pairs, such that increased oxygen fugacity will subdue Ti isotopic fractionation in global magmatic series. Similarly, we find negligible Fe isotopic fractionation in Rindjani bulk rocks and mineral separates, unlike previous studies. This is also likely due to the oxidized nature of the Rindjani differentiation suite, which leads to similar Fe3+/Fe2+ ratios in melt and minerals and decreases overall mineral-melt Fe fractionation factors. Paired Ti and Fe isotopic analyses may therefore represent a powerful tool to assess oxygen fugacity during differentiation, independent from Fe 3+ determinations of erupted samples. 

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3. A Continuum from Iron Oxide Copper-Gold to Iron Oxide-Apatite Deposits: Evidence from Fe and O Stable Isotopes and Trace Element Chemistry of Magnetite

   https://doi.org/10.5382/econgeo.4752  

   Rodriguez-Mustafa, Maria A.; Simon, Adam C.; del Real, Irene; Thompson, John F.H.; Bilenker, Laura D.; Barra, Fernando; Bindeman, Ilya; Cadwell, David (November 2020, Economic Geology)

   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. 

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4. Formation of iron oxide–apatite deposits

   https://doi.org/https://doi.org/10.1038/s43017-022-00335-3  

   Martin Reich, Adam C. (September 2022, Nature reviews)

   Renewed economic interest in iron oxide–apatite (IOA) deposits — containing tens to hundreds of millions of tonnes of Fe and substantial amounts of rare earth elements, P, Co and V — has emerged to supply the sustainable energy transition. However, the mechanisms that efficiently concentrate dense iron- rich minerals (for example, in ores up to ~90% magnetite) at the Earth’s near- surface are widely debated. In this Review, we discuss synergistic combinations of magmatic and hydrothermal iron- enrichment processes that can explain the available geochemical, petrological and geological IOA data. IOA deposits typically evolve from subductionrelated water- rich and chlorine- rich intermediate magmas under a wide temperature range, almost spanning the whole igneous–hydrothermal spectrum (from ~1,000 to 300 °C). Magmatic– hydrothermal fluids could efficiently scavenge Fe from magmas to form large IOA deposits (>100 million tonnes of Fe), whereas crystal fractionation and liquid immiscibility processes might account for more minor Fe mineralization occurrences. Igneous magnetite crystallization, volatile exsolution and highly focused transport of Fe- rich hydrothermal fluids through the crust under extensional tectonic conditions could be key factors enabling concentration of dense magnetite minerals in the less- dense upper crust. Future research should target both fertile and barren mafic–intermediate magmatic suites for distinctive signatures diagnostic of metallogenic fertility, to help unravel the genetic linkage between IOA and iron oxide–copper–gold systems. 

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5. Hydrothermal REE reaction paths in critical mineral deposits

   Gysi, A. P. (July 2021, Goldschmidt)

   Critical mineral deposits commonly form in magmatic-hydrothermal systems including carbonatites and/or alkaline syenites, and more evolved peralkaline granites where the rare earth element (REE) undergo a complex array of partitioning, transport and mineralization. Significant alteration and veining zones develop in these deposits and can be used to vector ore zones in the field [1]. The REE ore minerals typically reflect the characteristics of these systems, which are enriched in carbonate, fluoride, and phosphate or a combination thereof. The REE can also be incorporated into vein minerals such as calcite, fluorite and apatite where the REE3+ exchange for Ca2+ on the crystal lattice [2]. These minerals give us clues about the hydrothermal reaction paths of REE in critical mineral deposits. This study aims to: 1) present our recent findings from hydrothermal fluid-mineral REE partitioning experiments, 2) discuss thermodynamic models to simulate REE in critical mineral deposits, and 3) link the thermodynamic simulations to field observations. Hydrothermal fluid-calcite partitioning experiments were conducted between 100 and 200 °C by hydrothermal fluid mixing and precipitation [2] at near neutral to mildly alkaline pH (6 – 9). The REE concentrations in synthetic calcite crystals and aqueous fluids sampled in situ were used to fit the data to the lattice strain model [3] and using the Dual Thermodynamic approach [4]. A second type of experiment consisted of reacting natural fluorite and apatite crystals with acidic to mildly acidic (pH of 2 – 4) aqueous fluids in batch-type reactors to study the behavior of REE and mineral dissolution-precipitation reactions near crystal surfaces. The GEMS code package [5] was used to implement these new data into a thermodynamic model and simulate possible REE reaction paths in hydrothermal fluids. Two REE mineral deposits in New Mexico (Lemitar and Gallinas Mountains) present ideal case studies to illustrate how these models can be linked to field observations from natural systems. [1] Gysi et al. (2016), Econ. Geol. 111, 1241-1276; [2] Perry and Gysi (2020), Geochim. Cosmochim. Acta 286, 177-197; [3] Blundy and Wood (1994) Nature 372, 452-454; [4] Kulik (2006), Chem. Geol. 225, 189-212; [5] Kulik et al. (2013), Computat. Geosci. 17, 1-24. 

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