# Release of oxidizing fluids in subduction zones recorded by iron isotope zonation in garnet

Anna Gerrits, Edward Inglis, Besim Dragovic, Paul Starr, Ethan Baxter,

Kevin Burton

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Anna Gerrits, Edward Inglis, Besim Dragovic, Paul Starr, Ethan Baxter, et al.. Release of oxidizing fluids in subduction zones recorded by iron isotope zonation in garnet. Nature Geoscience, 2019, 12 (12), pp.1029-1033. 10.1038/s41561-019-0471-y. insu-02915506

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# Release of oxidizing fluids in subduction zones recorded by iron isotope zonation in garnet

Anna R. Gerrits<fnref n="1" />, Edward C. Inglis<fnref n="1" />,², Besim Dragovic<fnref n="1" />,³, Paul G. Starr<fnref n="1" />, Ethan F. Baxter<fnref n="1" />* and Kevin W. Burton⁴

Subduction zones are key regions of chemical and mass transfer between the Earth's surface and mantle. During subduction, oxidized material is carried into the mantle and large amounts of water are released due to the breakdown of hydrous minerals such as lawsonite. Dehydration accompanied by the release of oxidizing species may play a key role in controlling redox changes in the subducting slab and overlying mantle wedge. Here we present measurements of oxygen fugacity, using garnet-epidote oxybarometry, together with analyses of the stable iron isotope composition of zoned garnets from Sifnos, Greece. We find that the garnet interiors grew under relatively oxidized conditions whereas garnet rims record more reduced conditions. Garnet $\delta^{56}\text{Fe}$ increases from core to rim as the system becomes more reduced. Thermodynamic analysis shows that this change from relatively oxidized to more reduced conditions occurred during lawsonite dehydration. We conclude that the garnets maintain a record of progressive dehydration and that the residual mineral assemblages within the slab became more reduced during progressive subduction-zone dehydration. This is consistent with the hypothesis that lawsonite dehydration accompanied by the release of oxidizing species, such as sulfate, plays an important and measurable role in the global redox budget and contributes to sub-arc mantle oxidation in subduction zones.

Oxygen fugacity ($f_{O_2}$) is an intensive thermodynamic property that controls the speciation of multivalent elements, such as Fe, which regulates the ability of these elements to contribute to mineral-forming reactions<fnref n="1" />. Because of this, $f_{O_2}$ is an important chemical control in a given environment, such as a subduction zone, where the subducting slab experiences multiple changes in mineral assemblages as a result of changes in pressure-temperature (P-T) conditions. During subduction, components of the downgoing lithosphere, including sediments, altered mafic oceanic crust and serpentinized mantle, carry water and other volatiles into the subduction zone. It is widely accepted that, while there is major fluid loss during shallow subduction, hydrous phases in the downgoing slab, such as lawsonite, chlorite, chloritoid, phengite and serpentine, retain water to greater depths²⁻⁴. Lawsonite, CaAl₂Si₂O₇(OH)₂·H₂O, is a common hydrous mineral in mafic oceanic crust and contains ~12 wt% water²⁻⁴, making the metamorphic destabilization of lawsonite an important fluid-producing reaction within subduction zones. Much debate exists about the role of slab-derived fluids in oxidizing the overlying sub-arc mantle and how such fluids can be used to explain the oxidized and volatile-rich signatures observed in arc magmas relative to their mid-ocean-ridge basalt or ocean island basalt counterparts²,⁵⁻¹². The role of strongly oxidizing sulfur in governing subduction-zone redox processes has received much attention⁸,¹⁰,¹³,¹⁴, and recent work has suggested that the release of oxidizing sulfur species, such as sulfate ($SO_x$) species, may occur across the blueschist-eclogite facies transition, coincident with the breakdown of hydrous phases such as chlorite, talc and lawsonite¹⁴. Numerous studies suggest that the release of $SO_x$ species,

in fluids from the subducting slab, may play a key role in generating the elevated $f_{O_2}$ values proposed for the sub-arc mantle¹¹,¹³⁻¹⁹. Here we present a novel geochemical method on the scale of individual mineral growth zones to document the redox nature of the fluids released during key subduction dehydration events.

## Fe isotopes trace subduction-zone redox processes

To identify the source of these oxidizing (high-$f_{O_2}$) fluids, we use the exhumed rock record to constrain the $f_{O_2}$ of the residual slab mineral assemblage, which records redox changes resulting from dehydration reactions. There exist limited tools with which to measure the $f_{O_2}$ of the residual slab assemblages. Recent work has demonstrated the utility of garnet as a robust recorder of changing conditions during its formation and growth, including redox change²⁰ often spanning millions of years²¹⁻²³. Here we present the first high-precision stable Fe isotope measurements of individual garnet growth zones to examine the evolving $f_{O_2}$ of the downgoing slab during subduction. Changes in oxidation state are manifested in the $\text{Fe}^{2+}/\Sigma\text{Fe}$ ratio of minerals and result in a redox-driven fractionation of Fe isotopes, as heavier Fe isotopes (higher $\delta^{56}\text{Fe}$ values) are preferentially incorporated in bonds involving $\text{Fe}^{3+}$ relative to $\text{Fe}^{2+}$ when the system is in equilibrium²⁴,²⁵. Previous work on slab serpentinites has suggested that subducting slab material will lose isotopically light Fe with increasing grade, as a result of the release of fluids enriched in $\text{Fe}^{2+}$ in the form of aqueous $\text{Fe}^{2+}-\text{SO}_x$, $\text{Fe}^{2+}-\text{Cl}_2$ and $\text{Fe}^{2+}-\text{CO}_x$ complexes¹⁶. Despite this, no such effect has been resolvable on a whole-rock scale for metamafic lithologies within subduction zones²⁶. In our study, Fe isotope measurements are coupled with garnet-epidote oxygen barometry calculations²⁷,

<footnote n="1">¹Department of Earth and Environmental Sciences, Boston College, Chestnut Hill, MA, USA. ²Institut de Physique du Globe de Paris, Sorbonne Paris Cité, CNRS, Paris, France. ³School of the Earth, Ocean and Environment, University of South Carolina, Columbia, SC, USA. ⁴Department of Earth Sciences, Durham University, Science Labs, Durham, UK. *e-mail: ethan.baxter@bc.edu</footnote>
<figure><img src="image_2.png" /></figure>

**Fig. 1** | Backscattered electron imagery, and Fe isotope and oxybarometry data, for an analysed garnet from sample 09DSF-23E. a, Backscattered electron image of the garnet grain used for Fe isotope and oxybarometry analysis in sample 09DSF-23E. Epidote inclusions used in oxybarometry calculations are shown as red circles and sampling zones used for Fe isotope measurements are indicated by coloured regions. x, chemical center (core) of the garnet; amph, amphibole; ep, epidote; gt, garnet; qz, quartz; par, paragonite. b, Example of an epidote inclusion used in oxybarometry calculations. c, Example of a lawsonite pseudomorph that was avoided in oxybarometry calculations. d, δ⁵⁶Fe (top) and Δlog FMQ (bottom) values for sample 09DSF-23E plotted from garnet rim to rim. δ⁵⁶Fe values were analysed once for each garnet zone and have been plotted symmetrically to show trends from garnet rim to rim with measurements plotted as solid lines and associated errors indicated by the shaded regions. Error calculations for the δ⁵⁶Fe data are 2 s.d. of four isotopic analyses for each sample. Errors for the Δlog FMQ values are ±0.2 log units for the core (purple) and rim (red) points, calculated using the ±1kbar and ±40 °C errors associated with P-T estimates from thermodynamic modelling³⁰. The errors for the Δlog FMQ of the intermediate zones are calculated accounting for the slightly larger uncertainty in the P-T conditions and were calculated using upper and lower bounds for the possible P-T conditions (see further discussion in Methods and calculations in Supplementary Table 5).

using multiple epidote inclusions sampled radially within a single garnet grain (Fig. 1), to provide a record of $f_{O_2}$ change during garnet growth. This study proposes the use of the relationship between the oxidation state and the isotopic composition of Fe recorded at the mineral scale, within zoned garnet crystals, as a tracer of $f_{O_2}$ change and the release of fluids containing oxidizing species during metamorphic devolatilization of subducting oceanic crust.

## Fe isotope and $f_{O_2}$ variation in zoned garnets

Applying these methods, we present Fe isotope data and oxygen fugacity calculations (Figs. 1 and 2) for three zoned garnet crystals hosted within subducted and exhumed metabasalts from Sifnos, Greece. Plotting $f_{O_2}$ as Δlog FMQ (fayalite–magnetite–quartz) against δ⁵⁶Fe for samples 09DSF-23E and 09DSF-54A, illustrated in Fig. 2, shows a significant change from more oxidized garnet interiors (higher Δlog FMQ) with a lighter Fe isotope signature (lower δ⁵⁶Fe) isotopes to more reduced garnet rims (lower Δlog FMQ) with a higher δ⁵⁶Fe composition. Sample 06MSF-6C shows no significant change in calculated $f_{O_2}$ (Δlog FMQ) values from garnet interior to rim and only a minor variation in δ⁵⁶Fe, with lower δ⁵⁶Fe values in the garnet core than in the garnet rim (Fig. 2). Oxygen fugacity values calculated here are consistent with the typical range of subducted metabasaltic $f_{O_2}$ values²⁷⁻³⁰ in that they show more oxidized conditions, relative to the FMQ buffer, than unaltered mid-ocean-ridge basalt⁸. The data from the intermediate garnet zones of sample 09DSF-23E allow for the

possibility of a small increase in Δlog FMQ before the sharp drop to the rim. We can only speculate on the cause of this possible increase; perhaps it reflects the influx of an oxidizing fluid from a deeper dehydrating assemblage as it passes through the system. Regardless, the key result from our data is the sharp drop in $f_{O_2}$ from garnet interior zones to the rim: a drop that coincides with a sharp increase in δ⁵⁶Fe values.

These data show that there is significant redox change recorded in the mafic component of the slab during subduction at the mineral scale. To explore the possible link between these redox changes recorded in garnet and key dehydration reactions, thermodynamic modelling constraints on the P–T evolution of these samples are shown in Fig. 3 and Supplementary Figs. 6 and 7. For samples 09DSF-23E (Fig. 3) and 09DSF-54A (Supplementary Fig. 6), garnet growth spanned the onset of lawsonite breakdown along subduction-zone P–T paths for Sifnos, Greece²²,²³,³¹. To the contrary, garnets in sample 06MSF-6C probably ceased growing before lawsonite breakdown (Supplementary Fig. 7 and ref. ²³). This suggests that processes operating during the breakdown of lawsonite are responsible for causing the change in $f_{O_2}$ from garnet interior zones to the rim recorded in 09DSF-23E and 09DSF-54A. These observations support the hypothesis that the dehydration of lawsonite plays a key role in liberating fluids, which can then become carriers of oxidizing species, thus altering the redox state of the subducting metamorphic mineral assemblage. Overall, our data reveal that garnet interiors grew under relatively oxidized conditions, recording higher $f_{O_2}$ and lower δ⁵⁶Fe values, whereas
<figure><img src="image_3.png" /><figcaption>**Fig. 2 |** Fe isotope data, presented as δ<sup>56</sup>Fe values, plotted against the Δ log FMQ values for each zone (core, intermediate zones, rim).</figcaption></figure>

Each fO₂ data point represents a composite of all epidote inclusion-garnet pairs in that zone for the garnet. The arrows show the general trends in the data from the core (smallest circles) to intermediate zones (larger circles) to garnet rims (hexagons). Error calculations for the δ<sup>56</sup>Fe data are 2 s.d. of four isotopic analyses for each sample. All errors for Δ log FMQ are ±0.2 log units based on ±1kbar and ±40 °C error associated with P-T estimates from thermodynamic modelling<sup>30</sup>.

garnet rims record more reduced conditions with lower fO₂ and higher δ<sup>56</sup>Fe values, with this shift occurring during lawsonite dehydration.

Mechanisms for Fe isotope variation

While bulk rock Fe³⁺/ΣFe ratios can become decoupled from fO₂, the observed inverse relationship from garnet interiors to rims between increasing δ<sup>56</sup>Fe values and decreasing fO₂ (Fig. 1d) confirms that Fe isotope fractionation is linked to changing redox conditions. To account for the observed covariation between higher fO₂ and lower δ<sup>56</sup>Fe in garnet interiors relative to lower fO₂ and higher δ<sup>56</sup>Fe in garnet rims, we propose two possible mechanisms related to the release of oxidizing fluids. First, Fe isotope fractionation occurs as an open-system process, controlled by the solubility difference between Fe²⁺ and Fe³⁺. The higher solubility of Fe²⁺ relative to Fe³⁺ (ref. ³²) promotes Fe isotope fractionation, as lighter Fe isotopes are complexed into oxidizing fluids, in the form of Fe²⁺-SOₓ and Fe²⁺-COₓ complexes, which are subsequently removed from the system. The progressive removal of Fe²⁺-bearing, isotopically light fluids is recorded by prograde garnet growth in the residual source rock, and can explain the shift from light to heavy Fe isotope compositions between garnet interiors and rims. In this scenario, the Fe isotopes act as a tracer of the release of oxidizing species, with which Fe²⁺ preferentially complexes, within fluids released during the breakdown of hydrous mineralogy.

A second, complementary mechanism is that Fe isotope frac-
tionation is controlled by the bulk rock Fe³⁺/ΣFe ratio in a closed
system with respect to Fe, with changing oxygen fugacity. Fluids
charged with oxidizing species leave the system, reducing the fO₂
of the system, and leaving behind a more reduced residual min-
eral assemblage enriched in Fe²⁺. This fO₂ change does not alter the
whole-rock Fe isotope composition, but instead, as garnet more
readily incorporates isotopically light Fe²⁺, causes a shift towards

<figure><img src="image_4.png" /><figcaption>**Fig. 3 |** Pressure-temperature and mineralogical evolution of the Sifnos samples and model for the release of oxidizing fluids during subduction.</figcaption></figure>

a, P–T plot showing P–T path during garnet growth for samples 09DSF-23E (green) and 06MSF-6C (orange). The P–T path between the garnet core and rim for sample 09DSF-23E is based on Sifnos P–T paths by Dragovic et al.<sup>22,23</sup> and Groppo et al.<sup>31</sup>. The garnet zone symbols along the P–T path show the approximate P–T conditions at which these zones are interpreted to grow. Conditions of garnet growth for sample 09DSF-54A (not shown) are very similar to those for 09DSF-23E. Phases are labelled for fields showing P–T conditions where epidote, lawsonite and lawsonite + epidote are stable. Boundaries shown between these fields are schematic only as the bulk chemistry for each rock is different. b, Equilibrium cumulative modal mineral volume abundance (%) for sample 09DSF-23E along the prograde and retrograde P–T path. The onset of lawsonite breakdown releases a free fluid phase, shown as the light-blue field and marked by the white dashed line at -515 °C. c, Schematic diagram showing conditions for sample 09DSF-23E at labelled points 1, 2 and 3 on the P–T and cumulative modal abundance diagrams (a,b). Apparent minor differences between a and b reflect different methods of calculation related to the treatment of effective bulk chemistry; see Methods for discussion.
higher $\delta^{56}\text{Fe}$ values within garnet crystal rims due to simple Rayleigh fractionation. In this case, the Fe isotopes act as a more direct tracer of the changing bulk rock $\text{Fe}^{3+}/\text{\Delta}\text{Fe}$ ratio and $f_{\text{O}_2}$. While it is difficult to determine which is the dominant mechanism responsible for the observed Fe isotope fractionation, it is plausible that during garnet growth a combination of these two mechanisms could contribute, as both are driven by the release of oxidizing fluids during progressive dehydration.

## The release of oxidizing fluids during slab devolatilization

The large drop change in $f_{\text{O}_2}$ (Δlog FMQ) and increase in $\delta^{56}\text{Fe}$ values observed towards the garnet rim in samples 09DSF-23E (Figs. 1 and 2) and 09DSF-54A (Fig. 2) coincides with the interval of lawsonite breakdown and release of a free fluid phase as shown by thermodynamic modelling (Fig. 3 and Supplementary Fig. 6). Whilst the lawsonite breakdown reaction itself did not directly create the oxidizing nature of those fluids, recent work has shown that the sulfate mineral, anhydrite (CaSO$_4$), undergoes destabilization around the blueschist to eclogite transition at the same P–T conditions as the breakdown of lawsonite¹⁴. These fluids could be charged with $\text{SO}_x$ species for relatively oxidized rocks such as ours¹⁴. Other work has shown that the breakdown of sulfide minerals during subduction can also produce $\text{SO}_x$ species¹⁷. As a result, it is suggested that fluids produced during lawsonite dehydration could be charged with $\text{SO}_x$ complexes released from the coincident breakdown of sulfur-bearing phases, and thus serve as a powerful oxidant within these fluids. This hypothesis is fully supported by the trends in the Fe isotope and $f_{\text{O}_2}$ data, as light Fe is preferentially removed as $\text{Fe}^{2+}-\text{SO}_x$ complexes³², leaving a progressively more reduced residual assemblage. We conclude that these data support the hypothesis that the breakdown of lawsonite during subduction produces a fluid that influences the redox state and Fe isotopic signatures of the subducting mafic slab, which plays an important role in the overall global redox budget. In addition, this study provides evidence that Fe isotope variations in zoned garnets are a sensitive recorder of dehydration-driven redox change occurring during metamorphism of subducting oceanic lithosphere.

Much debate exists about the cause of the volatile enrichment and oxidized nature of arc magmas²,⁵⁻¹². It is widely accepted that fluids released from the subducting slab are responsible for altering the sub-arc mantle, the source region for arc magmas. A number of studies propose that these fluids, originating from dehydration of oceanic sediments, mafic crust and serpentinized mantle, are the agents of mantle wedge oxidation, and have shown that the transfer of oxidizing species from the slab to the sub-arc mantle within these metamorphic fluids is capable of causing significant shifts in the $f_{\text{O}_2}$ of the sub-arc mantle⁸,¹⁰,¹¹. Alternatively, it has been suggested that the oxidation of arc magmas may occur during magmatic differentiation or degassing¹²,¹⁵. This study provides evidence for redox change in the mafic component of the downgoing slab, inferred to be the result of the release of oxidizing fluids during lawsonite breakdown. By fingerprinting the source of oxidizing (high-$f_{\text{O}_2}$) fluids in the subducting slab, our study provides support for the idea that slab-derived fluids play an important role in the oxidation of the sub-arc mantle and related arc volcanic magmas.

## Online content

Any methods, additional references, Nature Research reporting summaries, source data, statements of code and data availability and associated accession codes are available at https://doi.org/10.1038/s41561-019-0471-y.

Received: 26 June 2019; Accepted: 12 September 2019; Published online: 28 October 2019

## References

1. Frost, B. R. in *Oxide Minerals: Petrologic and Magnetic Significance. Reviews in Mineralogy Vol. 25* (ed. Lindsley, D. H.) 1–9 (Mineralogical Society of America, 1991).

2. Schmidt, M. W. & Poli, S. Experimentally based water budgets for dehydrating slabs and consequences for arc magma generation. *Earth Planet. Sci. Lett.* **163**, 361–379 (1998).

3. Magni, V., Bouilhol, P. & van Hunen, J. Deep water recycling through time. *Geochem. Geophys. Geosyst.* **15**, 4203–4216 (2014).

4. Poli, S. & Schmidt, M. W. H₂O transport and release in subduction zones: experimental constraints on basaltic and andesitic systems. *J. Geophys. Res. Solid Earth* **100**, 22299–22314 (1995).

5. Tatsumi, Y. Formation of the volcanic front in subduction zones. *Geophys. Res. Lett.* **13**, 717–720 (1986).

6. Marschall, H. R. & Schumacher, J. C. Arc magmas sourced from mélange diapirs in subduction zones. *Nat. Geosci.* **5**, 862–867 (2012).

7. Breeding, C. M., Ague, J. J. & Bröcker, M. Fluid-metasedimentary rock interactions in subduction-zone mélange: implications for the chemical composition of arc magmas. *Geology* **32**, 1041–1044 (2004).

8. Kelley, K. A. & Cottrell, E. Water and the oxidation state of subduction zone magmas. *Science* **325**, 605–607 (2009).

9. Parkinson, I. J. & Arculus, R. J. The redox state of subduction zones: insights from arc-peridotites. *Chem. Geol.* **160**, 409–423 (1999).

10. Evans, K. A., Elburg, M. A. & Kamenetsky, V. S. Oxidation state of subarc mantle. *Geology* **40**, 783–786 (2012).

11. Debret, B. & Sverjensky, D. A. Highly oxidising fluids generated during serpentinite breakdown in subduction zones. *Sci. Rep.* **7**, 10351 (2017).

12. Lee, C. T. A. et al. The redox state of arc mantle using Zn/Fe systematics. *Nature* **468**, 681–685 (2010).

13. Pons, M.-L., Debret, B., Bouilhol, P., Delacour, A. & Williams, H. Zinc isotope evidence for sulfate-rich fluid transfer across subduction zones. *Nat. Commun.* **7**, 13794 (2016).

14. Tomkins, A. G. & Evans, K. A. Separate zones of sulfate and sulfide release from subducted mafic oceanic crust. *Earth Planet. Sci. Lett.* **428**, 73–83 (2015).

15. Evans, K. A. The redox budget of subduction zones. *Earth-Sci. Rev.* **113**, 11–32 (2012).

16. Debret, B. et al. Isotopic evidence for iron mobility during subduction. *Geology* **44**, 215–218 (2016).

17. Walters, J. B., Cruz-Uribe, A. M. & Marschall, H. R. Isotopic compositions of sulfides in exhumed high-pressure terranes: implications for sulfur cycling in subduction zones. *Geochem. Geophys. Geosyst.* **20**, 3347–3374 (2019).

18. Bénard, A. et al. Oxidising agents in sub-arc mantle melts link slab devolatilisation and arc magmas. *Nat. Commun.* **9**, 3500 (2018).

19. Canil, D. & Fellows, S. A. Sulphide–sulphate stability and melting in subducted sediment and its role in arc mantle redox and chalcophile cycling in space and time. *Earth Planet. Sci. Lett.* **470**, 73–86 (2017).

20. Ague, J. J., Baxter, E. F. & Eckert, J. O.Jr. High $f_{\text{O}_2}$ during sillimanite zone metamorphism of part of the Barrovian type locality, Glen Clova, Scotland. *J. Petrol.* **42**, 1301–1320 (2001).

21. Baxter, E. F. & Caddick, M. J. Garnet growth as a proxy for progressive subduction zone dehydration. *Geology* **41**, 643–646 (2013).

22. Dragovic, B., Baxter, E. F. & Caddick, M. J. Pulsed dehydration and garnet growth during subduction revealed by zoned garnet geochronology and thermodynamic modeling, Sifnos, Greece. *Earth Planet. Sci. Lett.* **413**, 111–122 (2015).

23. Dragovic, B., Samanta, L. M., Baxter, E. F. & Selverstone, J. Using garnet to constrain the duration and rate of water-releasing metamorphic reactions during subduction: an example from Sifnos, Greece. *Chem. Geol.* **314–317**, 9–22 (2012).

24. Polyakov, V. B. & Mineev, S. D. The use of Mossbauer spectroscopy in stable isotope geochemistry. *Geochim. Cosmochim. Acta* **64**, 849–865 (2000).

25. Schauble, E. A., Rossman, G. R. & Taylor, H. P. Theoretical estimates of equilibrium Fe isotope fractionations from vibrational spectroscopy. *Geochim. Cosmochim. Acta* **65**, 2487–2497 (2001).

26. Inglis, E. C. et al. The behavior of iron and zinc stable isotopes accompanying the subduction of mafic oceanic crust: a case study from Western Alpine ophiolites. *Geochem. Geophys. Geosyst.* **18**, 2562–2579 (2017).

27. Donohue, C. L. & Essene, E. J. An oxygen barometer with the assemblage garnet–epidote. *Earth Planet. Sci. Lett.* **181**, 459–472 (2000).

28. Boundy, T. M., Donohue, C. L., Essene, E. J., Mezger, K. & Austrheim, H. Discovery of eclogite facies carbonate rocks from the Lindás Nappe, Caledonides, Western Norway. *J. Metamorph. Geol.* **20**, 649–667 (2002).

29. Cao, Y., Song, S. G., Niu, Y. L., Jung, H. & Jin, Z. M. Variation of mineral composition, fabric and oxygen fugacity from massive to foliated eclogites during exhumation of subducted ocean crust in the North Qilian suture zone, NW China. *J. Metamorph. Geol.* **29**, 699–720 (2011).

30. Mattinson, C. G., Zhang, R. Y., Tsujimori, T. & Liou, J. G. Epidote-rich talc-kyanite-phengite eclogites, Sulu terrane, eastern China: P-T-fo2 estimates and the significance of the epidote-talc assemblage in eclogite. *Am. Mineral.* **89**, 1772–1783 (2004).
31. Groppo, C., Forster, M., Lister, G. & Compagnoni, R. Glaucophane schists and associated rocksfrom Sifnos (Cyclades, Greece): new constraints on the P-T evolution from oxidized systems. *Lithos* **109**, 254–273 (2009).

32. Hill, P. S., Schauble, E. A. & Young, E. D. Effects of changing solution chemistry on Fe³⁺/Fe²⁺ isotope fractionation in aqueous Fe–Cl solutions. *Geochim. Cosmochim. Acta* **74**, 6669–6689 (2010).

### Acknowledgements

E.F.B. acknowledges support from NSF grants EAR-0547999 for sample collection and OIA-1545903 for support during this project. A.R.G., B.D. and P.G.S. also acknowledge support from OIA-1545903. E.C.I. was supported as a postdoctoral research assistant on ERC starting grant PRISTINE: 637503 awarded to F. Moynier (IPG Paris).

### Author contributions

A.R.G. was responsible for the oxybarometry calculations, clean-lab preparation and analysis of Fe isotope compositions and wrote the manuscript. E.C.I. performed Fe isotope compositional analysis and contributed to method development. B.D. was responsible for P–T modelling. P.G.S. performed clean-lab preparation, oxybarometry calculations and

method development. E.F.B. and K.W.B. contributed by designing the project. All authors contributed to analysis and interpretation of data and editing of the manuscript.

### Competing interests

The authors declare no competing interests.

### Additional information

**Supplementary information** is available for this paper at https://doi.org/10.1038/s41561-019-0471-y.

**Correspondence and requests for materials** should be addressed to E.F.B.

**Peer review information** Primary Handling Editor(s): Melissa Plail; Heike Langenberg.

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© The Author(s), under exclusive licence to Springer Nature Limited 2019
## Methods

All three samples used for this study (09DSF-23E, 09DSF-54A and 06MSF-6C) are metasabaldes collected from Sifnos, Greece. GPS locations for each sample are as follows: 09DSF-23E, 37° 01.598' N, 24° 39.396' E; 09DSF-54A, 37° 00.930' N, 24° 39.360' E; 06MSF-6C, 37° 01.561' N, 24° 39.452' E.

**Oxygen fugacity calculations.** Oxygen fugacity methods and calculations are based on the oxygen barometer chemistry of Donohue and Essene³⁰ utilizing the THERMOCALC program and the thermodynamic database of Holland and Powell³³. To ensure that epidote inclusions are suitable for use in garnet-epidote oxygen barometer calculations, care was taken to identify primary epidotes, rather than compositionally patchy epidotes that reflect pseudomorphs after lawsonite. Epidote inclusions used in oxygenbarometry calculations (Supplementary Figs. 2–4) are in equilibrium with the surrounding garnet and show little to no cation zonation in backscattered electron analysis. Epidote inclusions interpreted to be lawsonite pseudomorphs show cation zonation visible in backscattered electron analysis and are accompanied by either paragonite or albite, representing breakdown products of lawsonite (Supplementary Fig. 5), and were avoided for this study.

Garnet and epidote endmember activities (Supplementary Table 3) were obtained at the appropriate P-T values for each epidote-garnet pair (Supplementary Tables 1 and 5–7) by inputting major element compositions of garnet and epidote minerals into the AX program (Tim Holland, University of Cambridge). Mineral compositions were acquired using wavelength dispersive spectrometry on a JEOL-JXA-8200 electron microprobe at the Massachusetts Institute of Technology (Supplementary Table 4). All spot analyses were carried out using an acceleration voltage of 15 kV, a current of 20 nA and an approximate spot size of 5 µm. Oxygen fugacity values (Supplementary Table 1) were calculated by inputting endmember activities of each epidote-garnet pair and the appropriate pressure and temperature into the thermodynamic modelling program THERMOCALC³³⁻³⁵ using the equation

$$2 \text{Ca}_2\text{FeAl}_2\text{Si}_3\text{O}_{12}(\text{OH}) \rightarrow 2 \text{Ca}_2\text{FeAl}_2\text{Si}_3\text{O}_{12} + \frac{\text{H}_2\text{O} + 1}{2} \text{O}_2 \quad (1) \\ \text{Epidote (P}_{533}\text{Czo}_{67}) \qquad \text{Garnet (Alm}_{33}\text{Grs}_{67}) \qquad \text{Fluid}$$

which can be rewritten to include grossular and almandine garnet endmembers:

$$12\text{Ca}_2\text{FeAl}_2\text{Si}_3\text{O}_{12}(\text{OH}) \rightarrow 8\text{Ca}_3\text{Al}_2\text{Si}_3\text{O}_{12} \quad (2) \\ \text{Epidote (P}_{533}\text{Czo}_{67}) \qquad \text{Grs Garnet} \\ + 4\text{Fe}_3\text{Al}_2\text{Si}_3\text{O}_{12} + 6\text{H}_2\text{O} + 3\text{O}_2 \quad \text{Alm Garnet} \qquad \text{Fluid}$$

Oxygen fugacity results are reported as Δ log FMQ, the difference between the calculated sample $f_{O_2}$ and the FMQ buffer at a given P-T:

$$\Delta \log \text{FMQ} = (\text{Sample } f_{O_2} P_{T} - \text{FMQ}_{P_{T}})$$

To obtain Δ log FMQ values for each $f_{O_2}$ value, the FMQ buffer was recalculated using the P-T conditions of each garnet-epidote oxybarometer pair used in $f_{O_2}$ calculations.

The calculation of the Fe³⁺ (andradite) component from the electron microprobe analyses was carried out using the charge-balance method described by Quinn et al.³⁶. As these charge-balance calculations of garnet compositions yield a low andradite component (<1.5% andradite component for garnet compositions in all three samples), garnet compositions used in oxygen fugacity calculations assume that all Fe in garnet is Fe³⁺. The effects of Fe³⁺ substitution within garnet is minor, with the exception of within highly andraditic garnets³⁷. The substitution of 15 mol% Fe³⁺/ΣFe, replacing Al in garnet, causes a shift of only ~0.1 log $f_{O_2}$ units²⁷. All Fe in epidote mineral formulas used in oxygen fugacity calculations is assumed to be Fe³⁺. Calculations assume a unit H₂O activity of 1 on the basis of low salinity measurements of garnet fluid inclusions from Cycladic metabasalts³⁷ and the low modal abundance of carbonate. Errors in Δ log FMQ (unless stated otherwise) are ±0.2 log units on the basis of the ±1 kbar and ±40 °C thermodynamic modelling error for P-T estimates presented by Palin et al.³⁸.

Oxygen fugacity calculations for epidote-garnet pairs in core and rim zones use estimated pressures and temperatures of formation for each sample based on garnet isopleth thermodynamic modelling for samples 09DSF-23E (Supplementary Fig. 7a,b), 09DSF-54A (Supplementary Fig. 7c,d) and 06MSF-6C (ref. ¹³). Without accurate P-T estimates for epidote inclusions in intermediate garnet zones, P-T values were estimated using a P-T path from Sifnos, Greece, with the prograde slope of Dragovic et al.²²,²³. Thermodynamic modelling constraints suggest that the majority of garnet growth, probably including all of the intermediate-zone growth, occurs at lower P-T conditions, shortly after the initial growth of the core (Fig. 3b and Supplementary Fig. 6). As a result, the P-T conditions chosen for the $f_{O_2}$ calculations within the intermediate zones are close to those of the core, projecting a small distance along the estimated P-T path (Fig. 3 and Supplementary Table 1). To estimate the uncertainties on these calculated intermediate-zone $f_{O_2}$ values, both the uncertainties in the P-T calculations (±1 kbar; ±40 °C) and the choice of P-T conditions were considered. Given that modelling shows very early garnet growth, it is plausible that all intermediate-zone garnet growth occurred at the core conditions. Thus $f_{O_2}$ values were calculated over a range of different P-T

conditions, considering that garnet growth may have occurred anywhere between the core conditions and the chosen P-T conditions along the P-T path. The minimum bound on the $f_{O_2}$ is calculated using the core conditions minus the P-T uncertainty (−1 kbar; −40 °C), whilst the maximum bound is calculated using the chosen P-T conditions plus the P-T uncertainty (+1 kbar; +40 °C) (Supplementary Tables 5–7). The resulting range of $f_{O_2}$ values is considered to represent the uncertainty in these intermediate-garnet-zone calculations, shown in Fig. 1c and Supplementary Fig. 1. In practice, the resulting uncertainty in the $f_{O_2}$ values is very similar to the ±0.2 log units used for the rim and core points (Fig. 1c and Supplementary Fig. 1).

**Fe isotopes.** Fe isotope ratios are reported as δ⁵⁶Fe and δ⁵⁷Fe using the IRMM-014 external standard with 2 s.d. reported error (Supplementary Table 2).

$$\delta^{56}\text{Fe} = ((\frac{^{56}\text{Fe}/^{54}\text{Fe}_{\text{sample}}}{^{56}\text{Fe}/^{54}\text{Fe}_{\text{IRMM-014}}}) - 1) \\ \delta^{57}\text{Fe} = ((\frac{^{57}\text{Fe}/^{54}\text{Fe}_{\text{sample}}}{^{57}\text{Fe}/^{54}\text{Fe}_{\text{IRMM-014}}}) - 1)$$

Garnet powders used in Fe isotope measurements were cleansed of inclusions and fully dissolved at Boston College (USA) before being passed through an Fe exchange chromatographic procedure and analysed for Fe isotope ratios at Durham University (UK). Growth zones were separated in individual garnet grains using the microdrilling techniques presented by Pollington and Baxter³⁹ to obtain three zones (core, zone 2 and rim) from samples 09DSF-54A and 06MSF-6C and four zones (core, zone 2, zone 3 and rim) from sample 09DSF-23E. Any visible inclusions were removed by handpicking and magnetic separation.

Samples were then put through a partial dissolution process, alternating dilute hydrofluoric and nitric acid steps to cleanse the garnet of inclusions. 10–50 mg of the picked garnet separate is heated at 120 °C and sonicated in a closed beaker with 1 ml deionized Milli-Q water and 5–90 µl concentrated hydrofluoric acid added to 1 ml of Milli-Q water on the basis of the starting amount of garnet for 120 min to dissolve inclusions. The residual garnet is then sonicated and heated at 120 °C for 120 min in 2 ml 7 M nitric acid to completely dissolve any secondary fluorides. This process is repeated until >50% of the original garnet has been dissolved. The inclusion-cleansed garnet residual is then fully dissolved using hydrofluoric acid, nitric acid and hydrochloric acid.

Tests were conducted to explore potential Fe isotope fractionation during the partial dissolution process described above. Pure gem-quality garnets with no visible inclusions from Mason Mountain Mine, North Carolina, were crushed and subjected to various partial dissolution procedures. To test potential fractionation due to the acids used, garnet was partially dissolved separately in 7 M nitric, concentrated hydrofluoric acid, and was also subjected to the full partial dissolution procedure described above. The resulting δ⁵⁶Fe compositions are all within error between pure garnet with no partial dissolution (0.015 ± 0.046), garnet after partial dissolution cleansing in nitric acid (0.085 ± 0.028), garnet after partial dissolution cleansing in hydrofluoric and nitric acid (0.08 ± 0.05) and garnet subjected to the full nitric–hydrofluoric multistage partial dissolution technique (0.06 ± 0.033). While the starting garnet may not have been perfectly pure, the partial dissolution cleansing removed those inclusions (with lower δ⁵⁶Fe) leaving behind a pure garnet with higher δ⁵⁶Fe. Importantly, after this initial stage of partial dissolution cleaning, all subsequent steps to further treat the garnets yielded identical δ⁵⁶Fe. This indicates the success of the method in removing the effects of non-garnet inclusions, but not altering the δ⁵⁶Fe of the pure garnet itself. On the basis of these results, the nitric–hydrofluoric acid multistage partial dissolution was deemed appropriate to cleanse garnet of inclusions without fractionating Fe isotope compositions. All garnet separates used in this study were subject to identical cleansing and preparation procedures.

The quantitative purification of Fe from the matrix elements was achieved using a protocol adapted from Dauphas et al.⁴⁰. In this method 1.2 ml of BioRad AG1-X8 (200–400 mesh) anion exchange resin was packed onto 11.5 ml total capacity polypropylene columns, and cleaned with passes of 10 ml of Milli-Q H₂O and 10 ml of 6 M hydrochloric acid repeated four times each. The resin was preconditioned with 2 ml of 6 M hydrochloric acid and the sample loaded onto the column in 250 µl of 6 M hydrochloric acid. The matrix was eluted from the retained Fe species by adding 8 ml of 6 M hydrochloric acid and discarded. Fe was quantitatively recovered from the column by adding 9 ml of 0.4 M hydrochloric acid, and subsequently collected into clean 15 ml Savillex Teflon beakers. The pure Fe solution was evaporated to dryness and brought back into solution in 2 ml of 0.5 M nitric acid before analysis by mass spectrometry. Previous calibration of this chromatographic ion exchange protocol demonstrated that the recovered Fe fraction was devoid of any isobaric elements (namely Cr and Ni) and totalled >99% of the Fe loaded into the column.

Fe isotope abundances were measured on a Thermo Scientific Neptune Plus MC-ICP-MS at Durham following the procedure of Weyer and Schwieters⁴¹. The instrument was run in medium-resolution mode, which gave a typical mass resolving power of ~6,500, adequate to discriminate between the ⁴⁰Ar¹⁴N⁺, ⁴⁰Ar¹⁶O⁺ and ⁴⁰Ar¹⁸OH⁺ polyatomic species that are isobaric on the ⁵⁴Fe⁺, ⁵⁶Fe⁺ and ⁵⁷Fe⁺ masses respectively. Samples were introduced to the plasma interface using an Elemental Scientific SIS quartz spray chamber coupled with a PFA 50 µl min⁻¹ nebulizer. All of the Fe masses were collected, as were ⁵³Cr⁺ and ⁶⁰Ni⁺, in the
movable Faraday collectors, mounted within the back end of the instrument.

$^{53}$Cr⁺ and $^{60}$Ni⁺ were used to correct, using the natural abundances, for any isobaric
interference from these elements on the $^{54}$Fe⁺ and $^{58}$Fe⁺ masses. In all cases this
correction had no effect on the calculated ratio, as Cr and Ni were quantitatively
removed from the sample solution before analysis by column chemistry.

Instrumental mass bias was corrected using standard sample bracketing, where IRMM-014 was used as the bracketing standard. Precision and accuracy were assessed by measuring both an in-house secondary reference solution (Durham FeWire) and an external geo-reference material (USGS BIR-1).

A total of 81 Durham FeWire analyses gave a mean of $\delta^{56}\text{Fe}$ of +0.23 ± 0.04‰ and a $\delta^{57}\text{Fe}$ of +0.37 ± 0.06‰ (n=81). Two aliquots of the BIR-1 geo-reference material were processed through two different chemistry batches and analysed a total of four times each. This gave an average $\delta^{56}\text{Fe}$ of +0.06 ± 0.028‰ and a $\delta^{57}\text{Fe}$ of +0.09 ± 0.021‰, which is in excellent agreement with published values for this standard¹²⁻¹⁴. The total procedural blank yielded <6 ng Fe, which is negligible when compared with the total amount of Fe processed through the columns. The Fe isotope data for the samples analysed as part of this study are reported in Supplementary Table 2 with an error of 2 s.d. of four isotopic analyses for each sample.

**Thermodynamic modelling.** To constrain the *P*−*T* conditions for garnet growth and the evolution of the metamorphic mineral assemblage during progressive subduction, *P*−*T* pseudosections and mineral modal plots were constructed using the thermodynamic program Perple_X (version 6.7.5)¹⁵ and the ‘ds 5.5’ update to the Holland and Powell¹³ internally consistent dataset. The chemical system MnO−Na₂O−CaO−K₂O−FeO−MgO−Al₂O₃−SiO₂−H₂O−TiO₂−Fe₂O₃ was used for all modelling. The following activity−composition models were used for phases involving solid solution: pyroxene and amphibole¹⁶, garnet⁷⁷, white mica⁴⁸,⁴⁹, chlorite⁵⁰, feldspar⁴¹, epidote and chloritoid⁴³, spinel⁵², ilmenite⁴³ and carbonates⁵⁴. In all the calculations, quartz, lawsonite, rutile, kyanite and sphene were assumed to be pure, with phase equilibrium calculations run in fluid-undersaturated conditions (see discussion below). Fe₂O₃ contents for the bulk compositions were evaluated by combining the average composition of phases with their respective volume abundances.

Bulk compositions used for all phase equilibrium calculations are shown in Supplementary Table 3, along with the $\text{Fe}^{2+}/\text{\Delta Fe}$ (by mole fraction) used for each sample. Whole-rock compositions were used in calculation of the *P*−*T* conditions of garnet growth initiation (garnet cores). As garnet is chemically zoned in both samples and the sequestration of components in zoned crystals can have a significant effect on both the effective composition of the rock and the resultant mineralogy⁵⁵, independent bulk compositions of rock matrices were obtained by physical separation of garnet crystals from a whole-rock volume. All bulk compositions (whole rocks and matrices) were determined by X-ray fluorescence spectroscopy using a Phillips 2404 X-ray fluorescence vacuum spectrometer at Franklin and Marshall College.

Path-dependent forward models take into account the continuous
fractionation of garnet and water, following Baxter and Caddick¹³, with a sequence
of regularly spaced *P*−*T* increments, where at each increment the composition and
modal abundance of all stable phases is predicted (Fig. 3b; Supplementary Fig. 6).
The models were run at 0.5 °C increments, with variable pressure increments.
The *P*−*T* paths used for this modelling (Fig. 3a) were chosen on the basis of the
individual *P*−*T* gradients derived from Dragovic et al.²², also utilizing additional
*P*−*T* constraints from Dragovic et al.²³. The whole-rock compositions were used
as the initial bulk compositions for the phase fractionation calculations. The
initial fluid contents for the phase fractionation calculations were determined on
the basis of repeat phase fractionation calculations to best model the observed
mineralogy (of garnet inclusions and matrix) and volumetric mineral abundances.
To reproduce the observed mineralogy, including the stable coexistence of
lawsonite and epidote during initial garnet growth, fluid-undersaturated
conditions were required. The initial fluid contents chosen for modelling of
samples 09DSF-23E, 09DSF-54A and 06MSF-6C were 6.0%, 4.0% and 2.0%,
respectively. The fluids in samples 09DSF-23E and 09DSF-54A were considered to
be fixed fluid H₂O−CO₂ compositions of 1 mol.% CO₂−99 mol.% H₂O (09DSF-23E)
and 10 mol.% CO₂−90 mol.% H₂O (09DSF-54A). The fluid in sample 06MSF-6C
was considered to be pure H₂O. These fluid compositions were estimated on the
basis of petrographic observations.

*P*−*T* pseudosections were calculated for the *P*−*T* ranges of 1.0−2.5 GPa and 400−650 °C using the same activity−composition models as listed above (Fig. 3a; Supplementary Fig. 7). The bulk composition used for the pseudosection modelling of each sample represents the effective bulk composition (garnet and water fractionated) calculated at 500 °C along the phase fractionation path. The fluid contents used for the pseudosections that estimate the *P*−*T* of garnet crystal cores are the same as those used as the initial fluid content for the phase fractionation calculations. For the pseudosections that estimate the *P*−*T* conditions of garnet crystal rims, the fluid content was based on predictions from the phase fractionation calculations for the remnant fluid content at the assumed *P*−*T* conditions of garnet rim growth (3.0% for 09DSF-23E and 2.7% for 09DSF-54A). An iterative analysis of fluid content resulted in broadly similar *P*−*T* pseudosections and predicted garnet rim *P*−*T* conditions.

**Rock descriptions.** Sample 09DSF-23E. A complete petrographic description of sample 09DSF-23E can be found in the work of Brooks et al.⁵³. Sample 09DSF-23E is a garnet–epitote blueschist that contains large (centimetre-sized), euheedral garnet crystals along with large (several-millimetre-sized) rhombohedral crystals of epidote. The remaining rock matrix consists of glaucophane, phengite, quartz, rutile, calcite, apatite, and trace albite and ilmenite. The cores of garnet crystals are inclusion rich, consisting of omphacite and quartz, with minor chlorite, phengite, paragonite, epidote, apatite, rutile and ilmenite. Garnet rims contain inclusions of glaucophane. Epidote inclusions in garnet occur in two varieties: (1) isolated subhedral inclusions that are regarded as primary epidote and (2) anhedral inclusions that are chemically zoned, always associated with albite and/or paragonite and interpreted to be pseudomorphic after lawsonite.

The above observations reflect the characteristics of a blueschist-overprinted eclogite, wherein the peak pressure assemblage consists of garnet, Na-pyroxene (omphacite), glaucophane, phengite, epidote, quartz, rutile and accessory phases. The blueschist overprint resulted in the growth of garnet rims, chemically zoned amphibole, rhombohedral epidote porphyroblasts and albite, and most notably the loss of matrix omphacite.

*Sample 09DSF-54A.* Sample 09DSF-54A is a garnet–epitote blueschist consisting of large, centimetre-sized garnet porphyroblasts and centimetre-sized rhombohedral epidote crystals. The rock matrix is dominated by glaucophane, with lesser amounts of phengite, Na-pyroxene (aegirine), quartz and titanite. Glaucophane and phengite define the foliation. Garnets are inclusion rich, comprised of Na-pyroxene (aegirine and jadeite), phengite, epidote, paragonite, quartz and rutile. As in sample 09DSF-23E, epidote occurs as a primary phase and in association with albite and/or paragonite, interpreted to be pseudomorphic after lawsonite.

Sample 09DSF-54A is regarded as a partially blueschist-overprinted eclogite. The peak eclogite assemblage consists of garnet, Na-pyroxene (aegirine and jadeite), phengite, epidote, quartz and rutile. The partial blueschist overprint is comprised of glaucophane, rhombohedral epidote porphyroblasts and lower-Si phengite, with the replacement of rutile by titanite.

*Sample 06MSF-6C.* A complete petrographic description of sample 06MSF-6C can be found in the work of Dragovic et al.²³. Sample 06MSF-6C is a garnet–epitote blueschist containing centimetre-sized porphyroblasts of garnet and millimetre-sized rhombohedral crystals of epidote crosscutting the matrix foliation. Garnet contains inclusions of quartz, Na-pyroxene (jadeite), glaucophane, chloritoid, paragonite, phengite and rutile. Unlike the above two samples, epidote was found as inclusions in garnet solely as pseudomorphic after lawsonite in garnet mantles and rims. Rare lawsonite was found as inclusions in garnet cores. The rock matrix is defined by glaucophane and white mica (both paragonite and phengite), but also consists of rhombohedral epidote, rutile, quartz and anhedral masses of jadeitic pyroxene.

Sample 06MSF-6C is interpreted to be a transitional blueschist to eclogite facies lithology. The peak assemblage is comprised of garnet, jadeite, glaucophane, paragonite, phengite, quartz and rutile. Matrix re-equilibration during the early stages of exhumation resulted in growth of rhombohedral epidote, paragonite and phengite that crosscuts the foliation.

**Data availability**

The data generated or analysed during this study are included in this published
article and its Supplementary Information files.

**References**

33. Holland, T. J. B. & Powell, R. An internally consistent thermodynamic data set for phases of petrological interest. *J. Metamorph. Geol.* **16**, 309–343 (1998).

34. Powell, R. & Holland, T. J. B. An internally consistent dataset with uncertainties and correlations: 3. Applications to geobarometry, worked examples and a computer program. *J. Metamorph. Geol.* **6**, 173–204 (1988).

35. Powell, R. & Holland, T. Optimal geothermometry and geobarometry. *Am. Mineral.* **79**, 120–133 (1994).

36. Quinn, R. J., Valley, J. W., Page, F. Z. & Fournelle, J. H. Accurate determination of ferric iron in garnets. *Am. Mineral.* **101**, 1704–1707 (2016).

37. Barr, H. Preliminary fluid inclusion studies in a high-grade blueschist terrain, Syros, Greece. *Mineral. Mag.* **54**, 159–168 (1990).

38. Palin, R. M., Weller, O. M., Waters, D. J. & Dyck, B. Quantifying geological uncertainty in metamorphic phase equilibria modelling; a Monte Carlo assessment and implications for tectonic interpretations. *Geosci. Front.* **7**, 591–607 (2016).

39. Pollington, A. D. & Baxter, E. F. High precision microsampling and preparation of zoned garnet porphyroblasts for Sm−Nd geochronology. *Chem. Geol.* **281**, 270–282 (2011).

40. Dauphas, N. et al. Chromatographic separation and multicollection-ICPMS analysis of iron. Investigating mass-dependent and-independent isotope effects. *Anal. Chem.* **76**, 5855–5863 (2004).

41. Weyer, S. & Schwieters, J. B. High precision Fe isotope measurements with high mass resolution MC-ICPMS. *Int. J. Mass Spectrom.* **226**, 355–368 (2003).
42. Millet, M. A., Baker, J. A. & Payne, C. E. Ultra-precise stable Fe isotope measurements by high resolution multiple-collector inductively coupled plasma mass spectrometry with a $^{57}$Fe-$^{58}$Fe double spike. *Chem. Geol.* **304**, 18–25 (2012).

43. Hibbert, K. E. J., Williams, H. M., Kerr, A. C. & Puchtel, I. S. Iron isotopes in ancient and modern komatiites: evidence in support of an oxidised mantle from Archean to present. *Earth Planet. Sci. Lett.* **321**, 198–207 (2012).

44. Sossi, P. A., Halverson, G. P., Nebel, O. & Eggins, S. M. Combined separation of Cu, Fe and Zn from rock matrices and improved analytical protocols for stable isotope determination. *Geostand. Geoanal. Res.* **39**, 129–149 (2015).

45. Connolly, J. A. D. The geodynamic equation of state: what and how. *Geochem. Geophys. Geosyst.* **10**, 10 (2009).

46. Diener, J. F. A. & Powell, R. Revised activity-composition models for clinopyroxene and amphibole. *J. Metamorph. Geol.* **30**, 131–142 (2012).

47. White, R. W., Powell, R. & Holland, T. J. B. Progress relating to calculation of partial melting equilibria for metapelites. *J. Metamorph. Geol.* **25**, 511–527 (2007).

48. Auzanneau, E., Schmidt, M. W., Vielzeuf, D. & Connolly, J. D. Titanium in phengite: a geobarometer for high temperature eclogites. *Contrib. Mineral. Petrol.* **159**, 1 (2010).

49. Coggan, R. & Holland, T. J. B. Mixing properties of phengitic micas and revised garnet-phengite thermometers. *J. Metamorph. Geol.* **20**, 683–696 (2002).

50. Powell, R., Holland, T. J. B. H. & Worley, B. Calculating phase diagrams involving solid solutions via non-linear equations, with examples using THERMOCALC. *J. Metamorph. Geol.* **16**, 577–588 (1998).

51. Fuhrman, M. L. & Lindsley, D. H. Ternary-feldspar modeling and thermometry. *Am. Mineral.* **73**, 201–215 (1988).

52. White, R. W., Powell, R. & Clarke, G. L. The interpretation of reaction textures in Fe-rich metapelitic granulites of the Musgrave Block, central Australia: constraints from mineral equilibria calculations in the system K₂O-FeO-MgO-Al₂O₃-SiO₂-H₂O-TiO₂-Fe₂O₃. *J. Metamorph. Geol.* **20**, 41–55 (2002).

53. White, R. W., Powell, R., Holland, T. J. B. & Worley, B. A. The effect of TiO₂ and Fe₂O₃ on metapelitic assemblages at greenschist and amphibolite facies conditions: mineral equilibria calculations in the system K₂O-FeO-MgO-Al₂O₃-SiO₂-H₂O-TiO₂-Fe₂O₃. *J. Metamorph. Geol.* **18**, 497–512 (2002).

54. Holland, T. & Powell, R. Activity-composition relations for phases in petrological calculations: an asymmetric multicomponent formulation. *Contrib. Mineral. Petrol.* **145**, 492–501 (2003).

55. Marmo, B. A., Clarke, G. L. & Powell, R. Fractionation of bulk rock composition due to porphyroblast growth: effects on eclogite facies mineral equilibria, Pam Peninsula, New Caledonia. *J. Metamorph. Geol.* **20**, 151–165 (2002).

56. Brooks, H. L., Dragovic, B., Lamadrid, H. M., Caddick, M. J. & Bodnar, R. Fluid capture during exhumation of subducted lithologies: a fluid inclusion study from Sifnos, Greece. *Lithos* **332–333**, 120–134 (2019).