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Mantle plumes are thermal upwellings that impinge on the lithosphere (French & Romanowicz, 2015; Montelli et al., 2006; Sleep, 1992). Capable of generating low degree mantle melts that more ably sample mantle heterogeneity, mantle plumes can produce ocean island basalts (OIB) (Dasgupta et al.; McKenzie & Onions, 1983; Stracke et al., 2005) and xenoliths wrenched from the lithospheric mantle (Frey & Roden, 1987). Both melts and xenoliths at ocean islands offer opportunities for oxybarometry; however, it remains challenging to interpret the relationship between lithospheric mantle peridotites and OIB. We begin with the volcanics.

Melt inclusions and submarine basalts erupting at and around the mantle plumes of Hawaii, Erebus, Iceland, and the Canary Islands record, on average, QFM +0.1 (7 XANES spectroscopic studies with n= 334 samples; Table 3.1, Fig. 3.2e). The Fe‐ XANES‐based fO₂ in this case is higher than the record of magnetite–ilmenite pairs (n=47) by 0.35 log units (t_statistic = 3.8, t_critical = 2, df = 69, p‐value = 0.003); however, we emphasize that these datasets have no samples in common. There is no meaningful difference between the fO₂s recorded by OIBs and MORBs. In the case of OIB volcanics, most XANES studies have either interrogated a geographic gradient in fO₂ (e.g., Shorttle et al., 2015) or the effect of differentiation (crystallization and degassing) on fO₂ (e.g. Brounce et al., 2017; Helz et al., 2017; Moussallam et al., 2016; Moussallam et al., 2014). We therefore observe a bimodal distribution of fO₂ recorded by plume‐affected glasses (Fig. 3.2e). Glasses erupted along mid‐ocean ridges that approach plumes, and glasses affected by degassing, record lower fO₂, while primitive and relatively undegassed melt inclusions record higher fO₂ (Fig. 3.2e, and see discussion). The authors of the detailed studies that have interrogated the fO₂ of plumes and plume‐affected ridges have inferred plume mantle source fO₂s anywhere from 0.4 to 2 log units higher than the average fO₂ recorded by the volcanics (filled circles in Figure 3.2e). In all cases, the authors have suggested that the mantle sources of these OIBs are more oxidized than those of MORB, and that this may be due to incorporation of recycled components (e.g. Brounce et al., 2017; Moussallam et al., 2019; Helz et al., 2017; Moussallam et al., 2016; Moussallam et al., 2014; Shorttle et al., 2015). Investigators have drawn such inferences by projecting the glass Fe³⁺/∑Fe ratios along compositional trends (e.g., to more primitive, less degassed, and more enriched compositions) or geographic trends (e.g., along a ridge toward a plume). We discuss some of these projections in greater detail in Section 3.4.1.1.

We have additionally compiled data from 13 studies to calculate the fO₂ of 143 ocean island xenoliths. Unlike all of the ridge peridotites and most of the arc peridotites, none of the published spinel compositions, save for Davis et al. (2017), were obtained using spinel standards with Fe³⁺/∑Fe ratios independently characterized. Without using standards to correct spinel Fe3+/∑Fe ratios, uncertainties in log fO₂ are roughly double those of samples calculated from corrected spinel Fe³⁺/∑Fe ratios (Davis et al., 2017, see Methods Appendix and Table 3.1). Despite, or because of, this limitation, we see that ocean island xenoliths record a wide range of fO₂, from QFM –2 to nearly QFM +4 with a mean equal to QFM +0.82 (±1.40). This is a much broader range, extending to lower fO₂, than compiled by Ballhaus (1993); the mean is within error of that compiled by Mallmann and O’Neill (2007), though again the variance is greater in the present compilation. Of the 143 xenoliths compiled here, 11 were identified by the original authors as pyroxenites (Sen, 1987; Tracy, 1980) which record a more oxidized mean fO₂ of QFM +1.44 (±0.63) than the whole set of OIB xenoliths (see red histogram overlain on Fig. 3.3d). The range of fO₂ recorded by OIB lavas falls within the range recorded by ocean island xenoliths; however, we cannot draw a genetic relationship between all lithospheric mantle xenoliths and the mantle melts that exhume them at plumes (compare Fig. 3.2e to 3.3d). Presumably, many of these xenoliths represent lithospheric mantle that has experienced limited chemical interaction with plume‐derived melts. Others are likely products of melt‐rock reaction between lithospheric peridotite and plume‐generated melts. The pyroxenite xenoliths were especially likely to have been derived in this way (e.g., Sen & Leeman, 1991), but it is unclear how many of the peridotites were also influenced by plume‐sourced melts. There is another caveat; we do not know the equilibration pressure at the closure temperature of these plume xenoliths. At ridges, we can reasonably infer pressure from peridotite thermometry because geothermal gradients are reasonably well‐characterized. That is not the case in plume settings, where thermal gradients are radial as well as vertical (e.g., Farnetani & Hofmann, 2010). In Figure 3.3e we demonstrate how equilibration at 2.5 GPa, instead of 0.6 GPa, would shift fO₂ down by two thirds of a log unit. This illustrates our community’s need for better peridotite mineral barometry.

We venture that a more appropriate comparison might be drawn between peridotitic ocean island xenoliths and ridge peridotites. Both may initiate as residues of melting at ridges, with the former transiting and cooling prior to interacting with a mantle plume. The range of fO₂s recorded by ocean island xenoliths encompasses the range of ridge peridotites but skews to higher fO₂s by 0.37 log units if we hold pressure constant at 0.6 GPa (t_stat = 2.6, t_crit = 2.0, df = 211, p‐value = 0.01). Notably, spinel‐olivine Fe‐Mg exchange (Li et al., 1995) records higher temperatures in the ocean island xenoliths than in the ridge peridotites. This difference in temperature could be the result of heating of the oceanic lithosphere beneath oceanic islands by the plume (Ballhaus, 1993), or it could result from exhumation of these xenoliths from greater depths than the depth of last equilibration experienced by ridge peridotites. In the former case, temperature‐dependent exchange reactions suggest that a mantle parcel preserving a record of hotter conditions should record lower fO₂ than a parcel at the same pressure that records cooler conditions (Birner et al., 2018). Consideration of these subsolidus reactions would thus predict ocean island xenoliths to be more reduced than ridge peridotites, in contrast to what we observe. If this interpretation is correct, then the difference in fO₂ between ridge peridotites and OIB xenolith source mantle prior to plume heating is even greater than the 0.6 GPa plots in Figure 3.3 suggest, perhaps driven by the interaction of some of these xenoliths with oxidized plume melts. In the latter case, changes in fO₂ due to changes in pressure would additionally have to be accounted for to make a direct comparison between ridge peridotites and OIB xenoliths. If no changes in mineral composition or mode are considered, the lower average fO₂ calculated assuming a higher pressure of equilibration (Fig. 3.3) suggests that if OIB xenoliths do generally sample deeper portions of the lithosphere than ridge peridotite, then average fO₂ of the two are comparable. While exchange reactions and modal changes during ascension and cooling may complicate this relationship, at this time, the data do not suggest that xenoliths recovered from plumes significantly differ in their fO₂ compared to peridotites recovered in the ridge setting. Constraints on pressure and effects of temperature on the fO₂ recorded by peridotites below their solidus remain poorly understood, and further work is needed to clarify the fO₂ signature of xenoliths entrained within plume lavas.