evidence employing redox proxies involving multivalent elements, such as V, Ce, and Eu, has been used to claim the constancy of the mantle redox state over the last 3.8 billion years, to values where carbonate melts are expected to be stable. As a consequence, it has been hypothesized that a mantle oxidation event occurred during the crystallization of the proto‐Earth from a magma ocean (Scaillet & Gaillard, 2011), which would imply that most of the deep carbon was mobilized to be released to the atmosphere. Basalts sampling the ambient mantle beneath spreading ridges () are subducted after 200–250 Ma. Older equivalents may be sampled from rare obducted ophiolites, Archean examples of which are, however, extremely rare and have a controversial tectonic setting, as also applies to mafic rocks occurring in greenstone belts (Bickle et al., 1994). One way to derive the redox state of Archean ambient mantle is to determine the V/Sc ratios and Fe3+/∑Fe of ancient basalts from the composition of well‐characterized Archean eclogitic rocks with spreading ridge‐derived mafic protoliths (Aulbach and Viljoen, 2015; Aulbach and Stagno, 2016; Aulbach et al., 2017; 2019). Although these studies find sympathetic behavior of both redox proxies and no evidence for a dominant control of metamorphism and partial melt extraction on V‐Sc‐Fe3+/∑Fe systematics, more work is needed to investigate the redox‐dependent partitioning of V and Fe3+ simulating the metamorphism of basaltic rocks coexisting with C‐O‐H fluids at high pressure and temperature. A complementary view of the deep convecting mantle sampled by thermo(chemical) plumes is provided by komatiites, which are hot, large‐volume melts emplaced predominantly in the Archean and Paleoproterozoic, enabling investigation of the partitioning of V between olivine and melts (Nicklas et al., 2018; 2019) and Fe3+ /∑Fe of melt inclusions contained in komatiite olivine (Berry et al., 2018).
Figure 2.2 LogfO2 (normalized to the FMQ buffer) calculated for mantle peridotite and eclogite xenoliths using oxy‐thermobarometers by Stagno et al. (2013; 2015; 2019). The blue and red lines are the oxygen fugacity calculated along a cratonic geotherm (~44 mW·m‐2) where diamonds (graphite) coexist with carbonatitic and kimberlitic melts, respectively. The purple line is the FeNi precipitation curve (Frost and McCammon, 2008), while the blue dashed line crosses the fO2 at which water is the main component of stable C–O–H fluids. At fO2 ≤ FMQ-3 log units fluids become increasingly CH4‐rich
(modified from Luth and Stachel, 2014).
2.3.2. Is the fO2 of the Transition Zone and Lower Mantle Recorded by Sublithospheric Diamonds?
Since no direct mantle rocks are available from depths greater than 200–250 km, the only tools to investigate the oxidation state at transition zone/lower mantle conditions and possible redox melting over time, in addition to the aforementioned mantle‐derived magmas, are experimental data combined with analyses of inclusions in natural sublithospheric diamonds. These diamonds, brought to the surface by kimberlitic eruptions, are often characterized by mineral inclusions representative of the transition zone and lower mantle (cpx, ringwoodite; majorite, NAL, Ca‐ferrite phase, bridgmanite, ferropericlase; Stachel et al., 2005; Pearson et al., 2014) that show geochemical evidence of the passage of oxidized C‐O‐H‐rich fluids from which they might have originated (Kaminsky, 2012; Kiseeva et al., 2018; Smith et al., 2016; Thompson et al., 2016; Walter et al., 2008). Recent discoveries of hydrous ringwoodite (Pearson et al., 2014) and ice‐VII (Tschauner et al., 2018) in diamonds from the transition zone further highlight the role of volatiles (other than C) bearing fluids in sublithospheric diamond formation. The origin of these fluids is often linked to subducted slabs widely accepted to be responsible for the oxygenation of the mantle from its initially reduced state in the wake of accretion and core formation (Lécuyer & Ricard, 1999). However, recent experimental investigations show that minerals like cpx, garnet, and majorite, which contain redox‐sensitive Fe and are the main constituents of metamorphosed oceanic crust, can incorporate large amounts of Fe3+ as result of increasing depth (i.e., pressure) and thus become significantly oxidized (Rohrbach et al., 2011; Kiseeva et al., 2018; Stagno et al., 2015, 2019). Such form of oxygen sequestration from minerals would result in reduced portions of the subducted slab, with diamond being the stable C form (i.e. redox freezing) and might explain the finding of large amounts of Fe3+ in majoritic inclusions trapped in diamonds (Kiseeva et al., 2018) as well as the occurrence of nanometer‐sized Mg‐ferrite inclusions in ferropericlase from superdeep diamonds (Harte et al., 1999; Wirth et al., 2014). While these natural observations open a scenario of a more oxidized rather than reduced Fe0‐saturated deep mantle as shown by thermodynamic and experimental predictions (Frost et al., 2004; Hu et al., 2019), some questions remain, such as the possibility that these inclusions do not reflect the whole‐mantle redox state but only local redox reactions.
2.3.3. Temporal and Spatial Evolution of the Redox State of the Asthenospheric Mantle
The oxygen fugacity of the ambient mantle, source of MORBs, and the deep convecting mantle sporadically sampled by plumes and kimberlitic melts both control the speciation of volatiles and hence the locus of the peridotite solidus and the nature of the melts that are produced (Wyllie & Huang, 1975; Canil & Scarfe, 1990; Green, 2015). In addition, mantle fO2 plays a role in mantle rheology and therefore dynamics, whereby more oxidizing conditions enhance concentrations and mobilities of lattice defects, resulting in a viscosity reduction (Cline et al., 2018). Given its importance, great efforts have been made to constrain mantle redox evolution, which is necessary to accurately capture mantle geodynamics and melting behaviour. For example, redox melting, whereby upwelling mantle will intersect the equilibrium C(graphite/diamond) + 2Fe2O3(melt) = 4FeO + CO2 (both in the melt) due to increasing fO2 with decreasing pressure, will occur in the depth interval of ~150 to 120 km at modern mantle redox conditions (Stagno et al., 2013), where MORB records fO2 of ~0.4±0.4 log units below the buffer (Frost & McCammon, 2008; Stagno, 2019). This process explains the production of kimberlites and some carbonatites, the rarity of which in the Archean has been ascribed to more reducing mantle conditions by 0.5 to 0.7 log units (Foley, 2011). Indeed, using the V/Sc redox proxy, a well‐characterized mantle eclogite suite from Lace in the Kaapvaal craton, with demonstrated spreading ridge‐derived origin, has yielded an fO2 (FMQ) of –1.7±1.1 for the ambient mantle at ca. 3 Ga (Aulbach & Viljoen, 2015). Based on this value, redox melting occurs at substantially shallower depths (~100 to 80 km; Aulbach & Stagno, 2016) with the formation of carbonate–silicate magmas, precluding the formation of carbonatite below typically thicker continental lithosphere, and possibly restricting the formation of kimberlites to unusually oxidized and/or warm mantle regions. Using additional mantle and orogenic eclogite suites sampling Proterozoic to Archean spreading ridges, the ambient mantle was shown to have evolved across the Archean–Proterozoic boundary, from ∆logƒO2 of ~FMQ‐1.19±0.33 to FMQ‐0.26±0.44 (Aulbach & Stagno, 2016). This evolution is exactly mirrored by that recorded in komatiites, which show an increase in ~1.3 log units between 3.48 and 1.87 Ga (Nicklas et al., 2018, 2019; Fig. 2.3a). This increase and the absolute fO2 values recorded across the Archean–Proterozoic are well within the range of, and therefore not in conflict with, proposed values after magma ocean solidification (~FMQ‐3 to FMQ) obtained from thermodynamic and climate modelling (Pahlevan et al., 2019). On the other hand, the compositions of inclusions in sublithospheric diamonds testify to the strong chemical and redox gradients, ranging from carbonates to solidified iron–nickel–carbon–sulfur melts (Brenker et al., 2007; Kaminsky et al., 2015; Smith et al., 2016). Indeed,
