# Potassium isotope evidence for slab-derived fluids in the sub-arc mantle

Kun Wang, Dmitri Ionov

► To cite this version:

Kun Wang, Dmitri Ionov. Potassium isotope evidence for slab-derived fluids in the sub-arc mantle. Earth and Planetary Science Letters, 2023, 619, pp.118315. 10.1016/j.epsl.2023.118315. hal-05265500

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# Potassium isotope evidence for slab-derived fluids in the sub-arc mantle

Kun Wang (王昆)¹,\* and Dmitri A. Ionov²

¹ Department of Earth and Planetary Sciences and McDonnell Center for the Space Sciences, Washington University in St. Louis, One Brookings Drive, St. Louis, MO 63130, USA

² Geosciences, Montpellier University, 34095 Montpellier, France

\* Correspondence author: wangkun@wustl.edu

Keywords: potassium isotopes; peridotite; subduction; mantle wedge; slab

Manuscript prepared for
**Earth and Planetary Science Letters**

Abstract length: 344

Main text: 7,309

References: 81

Tables: 2

Figures: 8

Supplementary Material: 1
Subduction recycles water, other volatiles and incompatible elements into Earth's mantle and plays a major role in exchanges between its surface and interior. Potassium (K), a highly fluid- mobile and incompatible alkali element, is strongly affected by these processes. The emerging K stable isotope system has been proposed as a new tracer of subduction events, e.g. to distinguish the inputs to arc magmas from sediments, altered basalts and other near-surface materials. Yet, it remains unknown if and how the K isotopic signatures of these components are affected by extraction of K-bearing media from dehydrating slabs and their transfer to the mantle wedge. Fragments (xenoliths) of sub-arc mantle carried to the surface by arc magmas may retain these slab-derived components. We report high-precision K isotope data on mantle harzburgite xenoliths from Avacha, an active arc volcano in Kamchatka (Russia) to further investigate the nature of slab-derived fluids and their impact on arc mantle. We find a narrow δ⁴¹K range for amphibole-rich cumulates from host andesite magmas (−0.53 ± 0.09‰ to −0.34 ± 0.09‰) that overlaps those of the bulk silicate Earth (BSE) and the upper continental crust (UCC). In contrast, the Avacha peridotites derived from the uppermost lithospheric mantle (refractory melting residues metasomatized by subduction zone fluids) show an extremely broad K isotopic range (δ⁴¹K ~ 3‰) and are commonly enriched in heavy K isotopes (up to +1.16 ± 0.08‰) relative to the BSE and UCC; some of the peridotites show the highest δ⁴¹K values for any reported mantle rocks. The new data on the Avacha xenoliths, along with available K isotope data on arc lavas and eclogites, suggest that slab-derived fluids may develop heavy K isotopic signatures, distinct from those in the original subducting materials (e.g., marine sediments and altered oceanic crust). These results provide direct evidence that (a) subduction dehydration and fluid percolation in the mantle lithosphere can significantly fractionate K isotopes, and (b) slab-derived fluids infiltrated the sub-arc mantle wedge. Due to the large K concentration difference between the depleted mantle and slab-derived fluids, K isotopes may be a sensitive tracer of fluid activity in sub-arc mantle.

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# 1. Introduction

Subduction zones play an important role in exchanges between Earth's surface and interior (Stern, 2002; Tatsumi and Kogiso, 2003). Through subduction, water, CO2 and other volatiles carried in sediments and altered oceanic crust (AOC) are recycled back into the mantle. While some of the volatiles are eventually transported to the deep mantle, large amounts of fluids are released to the mantle wedge to trigger metasomatism, partial melting and generate arc volcanism (Jarrard, 2003). These subduction-related fluids and melts are considerably enriched in fluid-mobile and incompatible components such as large-ion lithophile elements (LILE: K, Rb, Cs, Ba, light rare earth elements (LREE), etc.) (e.g., Hawkesworth et al., 1993; Spandler and Pirard, 2013).

Potassium isotopes have been recently proposed as a new tracer of sources and transport of fluid-mobile, incompatible elements in subduction zones and overlying mantle (Hu et al., 2020b, 2021a; Ionov and Wang, 2021; Liu et al., 2020; Sun et al., 2020). Both sediments and the AOC, two main recycled slab components, are enriched in K, while in contrast, the overlying melt-depleted lithospheric mantle is low in K, making K isotopes a potential sensitive tracer of slab-to-wedge transfers. The K2O concentrations [K2O] of global subducting sediments range from 0.15 to 3.93 wt.% (average 2.04 wt.%), mostly controlled by the ratio of K-rich detrital over K-poor biogenic phases (Plank and Langmuir, 1998). The δ41K values of these subducting sediments range from -1.31 to -0.02‰ (av. -0.48‰) (Hu et al., 2020b). This average is similar to the upper continental crust estimates (UCC: -0.44 ± 0.05‰; Huang et al., 2019), consistent with the view that K in sediments is mostly derived from detrital sources (Plank and Langmuir, 1998). In addition, authigenic clay (e.g., illite) formation during sediment diagenesis tends to take up K from seawater (average δ41K of seawater ~ +0.12 ± 0.07‰; Hille et al., 2019; Wang et al., 2020), but favors lighter K isotopes. The observed K isotopic range in subducting sediments reflects the mixing of detrital and diagenetic K. Older sediments show increasingly lighter K isotopic compositions (Hu et al., 2020b).

The AOC has significantly higher [K2O] (e.g., averages of 0.41 to 1.15 wt.% for ODP Site 801, IODP Sites U1365 and U1368; Hu et al., 2020b; Liu et al., 2021) than fresh MORB (mid-ocean ridge basalts; av. [K2O] ~0.19 wt.%; White and Klein, 2014) due to low-temperature hydrothermal alteration (e.g., Andrews, 1980; Hart, 1969) that forms minerals, such as saponite and celadonite, that incorporate K from seawater (Hu et al., 2020b; Liu et al., 2021). The AOC
also has a larger $\delta^{41}$K range (–0.76‰ to –0.17; Hu et al., 2020; Liu et al., 2021) than fresh MORB (–0.44 ± 0.17‰; Tuller-Ross et al., 2019a). The grand composite AOC at ODP Site 801 (–0.32‰; Hu et al., 2020b) is ~0.1‰ heavier than the bulk silicate Earth estimate (BSE: –0.43 ± 0.17‰; Tuller-Ross et al., 2019a) while the weighted average of AOC at IODP Sites U1365 and U1368 (–0.42‰; Liu et al., 2021) is indistinguishable from the BSE value.

Overall, the subducting sediments and the AOC may have slightly lighter or heavier, respectively, K isotopic compositions than the BSE allowing to trace their contributions in mantle-derived volcanic rocks (Hu et al., 2021a; Parendo et al., 2022; Sun et al., 2020). However, both the sediments and the AOC span a considerable $\delta^{41}$K range and largely overlap the BSE value. Thus, it is challenging to quantify relative contributions from different subduction components to arc magmas using the K isotope tracer alone. To further complicate the issue, the incompatible and fluid-mobile K is preferentially transferred to melts and fluids derived from sediments and AOC, and their isotopic compositions may evolve relative to their sources as this happens. Wang et al. (2022) found negligible K isotopic fractionation in high-pressure rocks formed from subducted marine sediments in a cold subduction zone and suggested that K isotopic signatures of sediments can be preserved during subduction zone metamorphism. In contrast, Liu et al. (2020) found a positive correlation between $\delta^{41}$K values and [K] in eclogites formed from subducted AOC, with the lightest K isotopic compositions (down to –1.64‰) in the most dehydrated K-poor rocks. They inferred that a complementary heavy-K-isotope enriched fluid (modeled $\delta^{41}$K range +0.23 to +1.04‰) was removed from the subducted AOC during dehydration and expelled to the overlaying mantle wedge.

Strongly melt-depleted, hence originally K-poor, harzburgites from the arc lithospheric mantle are appropriate samples for identifying such putative heavy-K-isotope signatures introduced by subduction fluids. Peridotite xenoliths may be more sensitive to such signatures than arc magmas, in which these heavy-K-isotope fluid components could be diluted by K scavenged from the fertile mantle wedge asthenosphere during fluid-induced melting (Hawkesworth et al., 1993). Yet, the K isotopic compositions of such direct samples from sub-arc mantle wedge have not yet been reliably constrained, hence the existence of such heavy-K-isotope fluids remains to be tested. To examine this hypothesis and gain new insights into the nature of the slab-derived fluids, as well as their role in sub-arc mantle metasomatism, we
obtained high-precision K isotope systematics of peridotite xenoliths and cumulates from their
host magmas from the Avacha volcano in the Kamchatka arc.

**2. Geological background and samples**

Avacha (Avachinsky) is an active stratovolcano in southern Kamchatka peninsula (53°15′N, 158°51′E), northeastern Russia (Fig. 1). It began erupting in late Pleistocene and is still one of the most active volcanos in Kamchatka. Avacha is part of the East Volcanic Front (EVF) in the Kuril-Kamchatka island-arc system, which formed due to the subduction of the Pacific plate under the Eurasian plate at the Kuril-Kamchatka trench. This subduction was initiated ~45 Myr ago, and the EVF was formed 5–7 Myr ago (see Bénard and Ionov, 2013; Ionov, 2010 for references). Avacha is ~200 km NE of the trench, where the subduction rate of Pacific plate is ~76 mm/year at a dipping angle of 53°; geophysical data show a low-velocity layer 70–120 km below the volcano, the Moho depth is ~37 km and the subducting slab is ~120 km below (Gorbatov et al., 1997; Levin et al., 2002).

While ultramafic xenoliths in andesitic arc volcanoes are generally rare and small, large and fresh mantle peridotites were found in basalt-andesitic ash and lapilli on the western slope of Avacha (Ionov, 2010). These xenoliths are spinel harzburgites, with high MgO (44–47 wt.%) and very low Al₂O₃ (0.4–0.9 wt.%), CaO (0.5–1.0 wt.%), TiO₂ (0.08–0.12 wt.%) and Na₂O (≤0.03 wt.%). Based on Al-Fe-Mg relations and heavy REE abundances, they formed by 28–35 wt.% melt extraction from fertile mantle at ≤1 to 2 GPa (Bénard et al., 2017a; Ionov, 2010) and represent the refractory lithospheric lid of the mantle wedge in a modern arc. The harzburgites contain 67–79 wt.% of olivine with Mg# of 0.907–0.917, 18–30 wt.% of orthopyroxene as well as accessory Cr-rich spinel (≤1 wt.%) and 1–3 wt.% of clinopyroxene exsolved from orthopyroxene on cooling. The mineral cores indicate equilibration temperatures ≤900–1000 °C. In addition, the harzburgites usually have trace interstitial amphibole (≤1 wt.%; <0.1mm in size; see Fig. 2C) and, less commonly, silicate glass at grain boundaries (often associated with empty vesicles, Fig. 2C) and in melt inclusions in spinel (Ionov, 2010; Ionov et al., 2011). These materials were formed from metasomatic media trapped in the sub-arc mantle (~37–55 km depth range based on and the absence of plagioclase and garnet) before the transport of the xenoliths to the surface. Analyses of grain rims and late-stage interstitial materials suggest their formation at
higher temperatures, oxidized conditions and high volatile abundances consistent with intrusion of slab-derived fluids (Bénard et al., 2018; Bénard and Ionov, 2013; Ionov, 2010).

The main K hosts in the harzburgites are interstitial amphibole (0.01–0.71 wt.% K₂O) and silicate glass (0.7–1.2 wt.% K₂O) (Bénard et al., 2017b; Ionov, 2010; Ionov et al., 2011). These late-stage metasomatic mantle materials contain oxidized components (e.g. sulfates), have high U/Th, Pb/Ce, Sr/Nd, δ³⁴S (+7 to +11‰), and hence formed from slab agents that contained dehydration products of serpentinised slab rocks in their sources (Bénard et al., 2018). Hopp and Ionov (2011) reported a range of slab-derived ‘atmospheric’ contributions to noble gas compositions in minerals separated from Avacha peridotites in this study that possibly reflect interaction of multi-component slab and mantle wedge fluids.

Apart from mantle peridotites, the Avacha volcanic rocks also carry rare magmatic cumulates consisting of coarse subhedral amphibole, minor olivine and silicate glass apparently deposited in intermediate andesitic magma chambers. The harzburgite xenoliths have mm-thick black coatings (selvages) of subhedral amphibole with interstitial silicate glass and olivine matching the modal and chemical compositions of the cumulates (Bénard and Ionov, 2013; Ionov, 2010). The selvages precipitated from host magma on the surface of the much colder (<1000 °C; Ionov, 2010) xenoliths. Overall, (a) both the selvages and cumulates are derived from the andesite magma, (b) the primary mineral assemblage of the harzburgites (coarse olivine, orthopyroxene and spinel) represents residues of old melt extraction that formed the mantle lithosphere, while (c) the late-stage amphibole and glass are products of recent slab-derived metasomatism. We want to emphasize here that the selvage amphibole and late-stage interstitial amphibole are very different in terms of location, size, and abundance. Selvage amphiboles wrap around the xenolith peridotites and crystallize from the host andesite lava, while the late-stage amphiboles are found within the xenolith and are located in interstitial pockets among olivines (see Fig. 2C). The selvage amphiboles are mm-sized, whereas the late-stage amphiboles are <0.1mm. Selvage amphiboles are the main minerals in the selvage, while late-stage amphiboles are present as accessory minerals (<1%) in the peridotites.

In this study, we analyzed fifteen whole-rock (WR) spinel harzburgite xenoliths and five samples from a cumulate and selvages from Avacha (Table 1). Two samples (Av-6 and Av-8) were previously analyzed and reported in Ionov and Wang (2021). The original peridotite xenoliths were >10 cm in size; their outer parts near contacts with host lava were removed (Fig.
2A). The xenolith cores were inspected visually and in thin sections to make sure they did not contain any materials related to host magma, crushed with hammer in plastic sheets, and ground in agate to yield ≥100 g of WR powders. Petrographic observations show no evidence for post-eruption alteration, e.g. olivine and interstitial glass are absolutely fresh (Ionov, 2010). Further, the WR powders show positive change of mass on ignition ranging from 0.53 wt.% to 0.65 wt.% and positively correlated with total iron as FeO (Bénard and Ionov, 2013; Ionov, 2010). The gain of mass on ignition means that oxidation of FeO to Fe₂O₃ is more significant than potential loss of volatiles (e.g., H₂O and CO₂ from alteration products) if they are present. Taken together, the petrographic and chemical data suggest no secondary alteration in the peridotites in this study.

About 1–2 g of bulk material from selvage on xenolith Av-18 (Bénard and Ionov, 2013) was ground to powder for analyses. Pure amphibole was handpicked from crushed magmatic cumulate Av-50 and selvages on Av-50 and peridotite xenoliths Av-18 and Av-39. These materials are referred to below as “host rock samples”. In contrast to the amphibole handpicked from the host rocks, separation and analyses of metasomatic amphibole and glass from the peridotite xenoliths are impossible because of their small size and low abundances (see Fig. 2C).

### 3. Methods

Aliquots of WR harzburgite powders (100–300 mg) and ~10 mg of host rock samples were digested in a concentrated HF-HNO₃ (3:1 v:v) acid mixture on a hot plate at 180°C for one week. Then, the samples were dried down and dissolved in a HCl-HNO₃ (3:1 v:v) acid mixture for two days, before being re-dissolved in 0.7 M HNO₃ for column chemistry. To ensure all K within the samples was dissolved, i.e. to avoid isotopic fractionation during digestion, additional aliquots of powders from four samples (Av-6, Av-8, Av-13, and Av-33; see **Table 1**) were digested using Parr high-pressure digestion vessels at 180°C in a Quincy Lab Model 10 lab oven for one week. After digestion, the samples were centrifuged and loaded onto K separation chromatography columns. A two-column K purification method after Chen et al. (2019) and Wang and Jacobsen (2016a) was applied. The columns were filled with 17 mL and 2.4 mL BioRad AG50-X8 100–200 mesh cation-exchange resin, respectively. Samples were eluted and K-cuts collected using 0.7 M or 0.5 M HNO₃ for the large and small columns respectively. The low-K harzburgite samples were eluted through the small columns two or three times to remove matrix elements, while the more K-rich host rock samples were eluted through the small columns only once. To
monitor column fractionation effects as well as accumulated column blank contamination, geological reference materials BHVO-2 and G-2 (Table 2) in different amounts (~0.01 mg, 0.1 mg, 10 mg) were passed through the columns alongside the harzburgites. Pre-cuts and post-cuts were collected for all samples at each column stage and analyzed with a Thermo Scientific iCAP Q quadrupole ICP-MS to monitor K loss in column chemistry. The final collected K-cuts for each sample were analyzed with the iCAP Q quadrupole ICP-MS to ensure the matrix element/K ratio was below 5 wt.%, which is the accepted threshold for generating measurable matrix effect on high-precision K isotopic analysis (Chen et al., 2019). The measured [K] concentrations in the final collected K-cuts are also reported in Table 1 with typical precision better than 10% (RSD). The total-procedure blank of our method is ~13 ng K (Liu et al., 2020), which is negligible, even for the lowest-K xenolith sample.

The K isotopic compositions of the peridotites, host rock and geological reference materials were measured using a Thermo Scientific Neptune Plus MC-ICP-MS as in Chen et al. (2019). The same analytical method was also used in Ionov and Wang (2021) for similarly low-K peridotite xenoliths. Samples were introduced into the MC-ICP-MS with an Elemental Scientific Apex Omega high sensitivity desolvating nebulizer to increase signal intensity and reduce the Ar hydride interference. Each sample was measured ~10 times (Table 1). The K isotopic compositions are expressed using the conventional $\delta^{41}K$ notation, where $\delta^{41}K = [(^{41}K/^{39}K)_{sample}/(^{41}K/^{39}K)_{NIST SRM3141a}-1] \times 1000$.

The average K isotopic compositions of peridotite and host-rock samples and their uncertainties (95% confidence interval; 95% c.i.) are reported in Table 1. The long-term (20 month) reproducibility (2 s.d.) of this method was evaluated as ±0.11‰ (Chen et al., 2019). Four peridotite samples (Av-6, Av-8, Av-13, Av-33) were replicated, i.e. digested (in Parr high-pressure vessels), chemically processed, and analyzed independently to monitor full-procedure reproducibility (Table 1). Values for reference materials BHVO-2 and G-2 processed and analyzed in the same sessions (Table 2) along with the samples agree well with literature data (e.g., Chen et al., 2019). BHVO-2#1 was analyzed daily for two weeks during our xenolith analysis sessions, with 2s.d. of ±0.12‰, which is consistent with our long-term reproducibility of ±0.11‰ (Chen et al., 2019).

## 4. Results
The K concentrations [K] in the Avacha harzburgite xenoliths in this study (Table 1) range from 6 to 32 ppm (av. 16 ppm, n = 15). The [K] values for four samples digested in Parr high-pressure vessels agree well with those for replicated samples digested on a hotplate (19 vs. 17 ppm; 27 vs. 21 ppm; 13 vs. 17 ppm; 27 vs. 32 ppm). The absence of systematic [K] differences between the two digestion methods suggests that both methods completely dissolve K-bearing phases in the xenoliths. The low [K] are consistent with X-ray fluorescence analyses of these samples (Ionov, 2010), which yielded values below detection limits (<0.01 wt.% K₂O or <83 ppm K). The main K carrier in these peridotites is amphibole which typically contains 0.2–0.3 wt.% K₂O, while much less common silicate glass has 0.13–1.1 wt.% K₂O (Bénard et al., 2016; Ionov et al., 2011). The whole-rock [K] obtained in this study are roughly correlated with amphibole modal abundances, e.g. [K] estimates for peridotites containing 0.4–1.0 wt.% amphibole with 0.25 wt.% K₂O range from 8 to 21 ppm K (Ionov, 2010; Supplementary Data Table S1). Abundant K-rich (0.7-1.2 wt.% K₂O) glass inclusions in spinel explain higher [K] of ~30 ppm in Av-33. The [K] in amphibole from the selvages and cumulate (1521–1992 ppm, i.e. 0.18–0.24 wt.% K₂O) is two orders of magnitude higher than in the WR peridotites (8–32 ppm), but similar to interstitial amphibole in the peridotites (0.01–0.71 wt.% K₂O, averaging at 0.25 wt.%; EPMA data from Ionov, 2010).

The δ⁴¹K values in the 15 harzburgite xenoliths range from -1.70 ± 0.19‰ to +1.16 ± 0.08‰ (Fig. 3). The four samples replicated using different digestion methods show similar δ⁴¹K values generally within our long-term analytical reproducibility; small discrepancies can be attributed to the nugget effect of K-bearing phases. This ~3‰ δ⁴¹K range for the Avacha peridotites is much larger than those reported for unaltered MORB (-0.25‰ to -0.66‰, av. -0.44 ± 0.17‰) (Tuller-Ross et al., 2019a) and continental upper crustal rocks (diorite, granodiorite, granite, loess, shale, graywacke, pelite, tillite) with δ⁴¹K of -0.68‰ to -0.12‰ (Huang et al., 2020). The ~3‰ δ⁴¹K range for the Avacha peridotites is also larger than the K isotopic range of ~2‰ (-2.11±0.06‰ to -0.06±0.06‰) for global whole-rock peridotite xenoliths (excluding two Avacha samples) reported by Ionov and Wang (2021). Most notably, many Avacha xenoliths are significantly enriched in heavy K isotopes compared to any other peridotite xenoliths analyzed so far (Fig. 3).

In contrast to the xenoliths, the five host-rock samples from Avacha show a restricted δ⁴¹K range, from -0.53 ± 0.09‰ to -0.34 ± 0.09‰, which is typical for crustal rocks (Table 1) and overlaps estimates for the BSE (-0.43 ± 0.17‰; Tuller-Ross et al., 2019a) and the UCC
($-0.44 \pm 0.05\%$; Huang et al., 2020). Thus, the amphibole-rich selvages that enclose the Avacha xenoliths differ significantly in K isotopic composition from the Avacha peridotites. This further suggests, in addition to the petrographic data and extremely low [K] in WR Avacha peridotites, their negligible contamination from host magma. Considering that [K] in the hosts is two orders of magnitude higher than in the peridotites, any contamination from such hosts would dominate peridotite K budget and push the measured K isotopic composition of the peridotites to the BSE value ($\sim-0.5\%$). By contrast, our data show a very broad $\delta^{41}$K range and no clustering near the BSE value (Fig. 3).

# 5. Discussion

## 5.1. The K isotopic composition of the lithospheric mantle estimated from peridotites

Peridotite xenoliths have been commonly used to estimate the chemical composition of the depleted MORB mantle (e.g., Carlson and Ionov, 2019) and the BSE (e.g., McDonough and Sun, 1995; Palme and Nickel, 1985) and, of late, to constrain their stable isotope composition (e.g., Bindeman et al., 2022; Teng, 2017). Ionov and Wang (2021) recently provided $\delta^{41}$K for over 60 geographically diverse mantle xenoliths and their minerals. They argued that the average $\delta^{41}$K for least altered and least metasomatized fertile lherzolites from their suite ($-0.57 \pm 0.28\%$) provides a sound BSE estimate, which agrees with the averages for global unaltered oceanic basalts ($-0.43 \pm 0.17\%$; Tuller-Ross et al., 2019a), Hawaiian basalts and a weighted average of continental and oceanic igneous rocks ($-0.43 \pm 0.10\%$; Hu et al., 2021b).

By contrast, veined and strongly metasomatized mantle peridotites show a much larger $\delta^{41}$K range than the fertile lherzolites that is not typically seen in crustal igneous rocks (Fig. 3). Ionov and Wang (2021) reported a $\sim2\%$ $\delta^{41}$K range ($-2.11 \pm 0.06\%$ to $-0.06 \pm 0.06\%$) for mantle xenoliths from the Siberian craton and off-craton localities in central and eastern Asia, Germany and Australia, that is nearly an order of magnitude greater than for unaltered oceanic basalts ($<0.4\%$ overall and $<0.2\%$ for most of them) (e.g., Hu et al., 2021b; Tuller-Ross et al., 2019a). The latter values suggest that high-temperature magmatic processes such as partial melting, fractional crystallization cause negligible or very small K isotopic fractionation (except in high-silica rocks; see examples in Huang et al., 2023; Wang et al., 2023). Thus, the large $\delta^{41}$K range reported for the peridotites should rather be attributed to kinetic isotope fractionation (e.g., Watkins et al., 2017) related to element diffusion during mantle metasomatism caused by magma
or fluid percolation (Ionov and Wang, 2021). Kinetic isotopic fractionation in the mantle has been previously invoked for many other isotopic systems, e.g. Li, Mg, Fe, Ca, Cr (see reviews in Teng et al., 2017 and references therein).

The data in this study show that positive δ⁴¹K values (0.40 ± 0.13‰, 0.62 ± 0.13‰) for two Avacha peridotites reported in Ionov and Wang (2021) are not an aberration, but typical of the mantle beneath Avacha (8 out of 15 peridotites in this study with an average of 0.51 ± 0.34‰, 1 s.d.). The Avacha peridotites show the highest δ⁴¹K values (up to +1.16 ± 0.08‰) reported for any mantle-derived rock, i.e. at least ~1‰ higher than any xenoliths outside subduction zones (≤-0.06 ± 0.06‰; Ionov and Wang, 2021) (Fig. 3). Further, we confirm that the K isotopic composition of the lithospheric mantle is highly heterogeneous both locally (-1.70 ± 0.19‰ to +1.16 ± 0.08‰ for Avacha) and globally (-2.11 ± 0.06‰ to +1.16 ± 0.08‰ for mantle samples in this study and Ionov and Wang (2021)). While the Avacha peridotites are the only subduction zone mantle samples analyzed so far for K isotopes, they show a greater δ⁴¹K range than the xenoliths outside subduction zones (~3‰ vs. ~2‰). Possible causes for the characteristic heavy K isotope enrichments in subduction zone peridotites are explored below.

**5.2. Mechanisms for K isotopic fractionation in the lithospheric mantle wedge**

The Avacha peridotites reported here were initially formed as residues after high-degree melt extraction (Section 2) and, as a result, are strongly depleted in K (down to 6 ppm) relative to fertile mantle estimates (240 ppm; McDonough and Sun, 1995) because K is highly incompatible during partial melting. Studies on volcanic rocks from Hawaii and Iceland (Hu et al., 2021b; Tuller-Ross et al., 2019b) show that high-temperature magmatic processes cannot generate large (>0.4 ‰) K isotopic fractionation. Further, the Avacha harzburgites show an extremely broad δ⁴¹K range, but narrow major oxide ranges indicating similar melting degrees (28–35 wt.%; Bénard et al., 2017; Ionov, 2010). Thus, melt extraction can be ruled out as the mechanism to yield the ~3‰ K isotopic fractionation among Avacha peridotites.

It may not be a coincidence that the only three peridotites with δ⁴¹K values within the ranges for the BSE and igneous rocks are those with the lowest [K] of 6–8 ppm: Av-12 (-0.49 ± 0.07), Av-15 (-0.53 ± 0.06) and Av-G-2 (-0.47 ± 0.09) (Fig. 4A and 4B). These samples may have remained unaffected by post-melting K inputs and thus retained the δ⁴¹K record of the initial
residual mantle. In such a case, the higher [K] in the other peridotites (8–32 ppm) indicate post-melting K additions that may also be responsible for their K-isotope variations.

In summary, the large $\delta^{41}$K range (–1.70 ± 0.19‰ to +1.16 ± 0.08‰) in the Avacha peridotites, the only lithospheric arc mantle samples yet analyzed, cannot be explained by initial melt extraction (nor post-eruption alteration or host-lava contamination). We will explore the following two mechanisms to explain the K isotopic range in these samples, both implying metasomatic K additions: 1) kinetic isotopic fractionation due to K diffusion; and 2) metasomatism by isotopically distinct fluids associated with the slab dehydration.

### 5.2.1. Diffusion and kinetic isotopic fractionation

Large isotopic fractionations in mantle rocks have been found for many elements, like Li, another alkali metal, with kinetic diffusion effects playing a major role (e.g., Jeffcoate et al., 2007; Lundstrom et al., 2005). Lithium has very high diffusivity in silicate melts and minerals because of its small ionic radius and low valence. Faster diffusion of $^{6}$Li than $^{7}$Li leads to large isotopic fractionations, as demonstrated by laboratory experiments (Dohmen et al., 2010; Holycross et al., 2018; Richter et al., 2003, 2014b, 2017). Two mechanisms were proposed for such processes in peridotites (e.g., Penniston-Dorland et al., 2017): (1) Li diffusion into minerals from a Li-rich grain boundary melt or fluid, and (2) Li redistribution between minerals due to changes in (temperature-dependent) inter-mineral Li partition coefficients during slow cooling. The latter process, however, can only explain inter-mineral fractionation, but not isotopic differences between bulk rocks (e.g., Penniston-Dorland et al., 2017).

Potassium has a larger ionic radius, hence lower diffusivity than Li (Richter et al., 2003). Yet, significant (>10‰) K isotopic fractionation was attained in laboratory diffusion experiments (Richter et al., 2014a; Zhang, 2022). The diffusion front becomes $^{39}$K enriched (i.e., develops lower $\delta^{41}$K values) because $^{39}$K diffuses faster than $^{41}$K. Ionov and Wang (2021) reported $\delta^{41}$K data on composite mantle xenoliths (spinel lherzolites cross-cut by phlogopite-rich veins) that are consistent with kinetic isotope fractionation by diffusion from fluid channels (later solidified as veins) to host rocks. They found consistently lower $\delta^{41}$K in the host peridotites than in the veins (i.e., excess $^{39}$K in the hosts due to its faster diffusion from fluid channels).

We posit that peridotites situated close to melt or fluid percolation front were enriched in lighter K isotopes. According to mass-balance, the remaining melt or fluid became increasingly
enriched in heavier K isotopes, similar to what was inferred for composite mantle xenoliths reported by Ionov and Wang (2021), but possibly on a larger scale. Overall, such diffusive mechanisms may be responsible for the large K isotopic variations observed in worldwide mantle xenoliths including some Avacha peridotites in this study (Fig. 3).

To illustrate this diffusive K isotopic fractionation mechanism in peridotites, we follow Zhang and Gan (2022) to calculate a 1D model of K isotopic fractionation during diffusion from a K-rich metasomatic melt/fluid to a K-poor peridotite.

The diffusion coefficient of K in olivine has not been determined. Therefore, we will establish an upper bound by using the diffusion coefficient of K in melts (referred to as $D_{max}$), and a lower bound by using the diffusion coefficient of K in other minerals (referred to as $D_{min}$). According to Ni (2012), the empirical equation for the diffusion coefficient of K in basaltic melts is as follows (with units of m²/s):

$$ \ln D_{max} = -14.81 - \frac{11125}{T} \quad (1) $$

The empirical equation for the diffusion coefficient ($D_{min}$; unit: m²/s) of K in albite is (Giletti and Shanahan, 1997):

$$ D_{min} = 3.98 \times 10^{-4} \cdot e^{-\frac{35603}{T}} \quad (2) $$

This diffusion coefficient of K in albite is comparable to those of trace elements (e.g., Na and Li; no K data) in olivine at equal temperature (i.e., a single experiment at 1,300 °C; Spandler and O'Neill, 2010). The diffusion in peridotites from a metasomatic melt/fluid is likely through grain boundaries and microchannels (Jeffcoate et al., 2007; Rudnick and Ionov, 2007). Experiments show that the effective diffusion of H in olivine including diffusion through grain boundaries could be orders of magnitude faster than the olivine lattice diffusion alone (Demouchy, 2010). Thus, the diffusion coefficient of K in albite is considered here as the lower bound. The effective diffusion coefficient of K in peridotites is likely to be between $D_{max}$ and $D_{min}$.

In both equations, T is temperature (K). Here we use 900 °C based on the estimated trapping temperature of the low-T melt inclusions in Avacha peridotites that likely formed due to the percolation of hydrous fluids or low-T melts in the arc mantle lithosphere (Ionov et al., 2011). We also calculated for comparison diffusion models at other temperatures (see Supplementary
Fig. S1 for T range from 900 °C to 1200 °C; Note that the temperature-dependent diffusion coefficient, D, and time scale, t, are trade-offs; see Eq. 5 below).

The relationship between the diffusion coefficients of isotopes ³⁹K and ³⁹K is mass-dependent and can be described using this empirical equation (Richter et al., 1999):

$$ \frac{D_{41}}{D_{39}} = \left( \frac{m_{39}}{m_{41}} \right)^{\beta} $$

(3)

where m is the isotope mass and β is an empirical fit parameter. The β values for K isotope diffusion in basaltic melts were experimentally determined to range from 0.090 to 0.116 at 1260–1500 °C (Zhang, 2022). Here we use 0.1 for β.

As ³⁹K is the most abundant isotope of K, we assume:

$$ D_{39} = D_{K} $$

(4)

The analytical solution for the K concentration (C) at x (distance to the boundary) and t (time) is (Carslaw and Jaeger, 1959; Crank, 1975):

$$ C = 0.5(C_{Low} + C_{High}) + 0.5(C_{High} - C_{Low})\erf\frac{x-x_0}{\sqrt{4Dt}} $$

(5)

where $x_0$ is the boundary location, $C_{Low}$ is the K concentration in the peridotite, and $C_{High}$ is the K concentration in the metasomatic melt/fluid. We use $C_{Low} = 6$ ppm (i.e., refractory peridotites) and $C_{High} = 60, 300, \text{ or } 600$ ppm (i.e., hypothetic metasomatic melt/fluid).

The K isotopic fractionation at x and t relative to the starting composition (BSE for both peridotite and metasomatic melt/fluid) is (Zhang and Gan, 2022):

$$ \Delta^{41}K = \left( \frac{\left( 1 + \frac{C_{High}}{C_{Low}} \right) + \left( \frac{C_{High}}{C_{Low}} - 1 \right) \erf \frac{x-x_0}{\sqrt{4D_{41}t}}}{\left( 1 + \frac{C_{High}}{C_{Low}} \right) + \left( \frac{C_{High}}{C_{Low}} - 1 \right) \erf \frac{x-x_0}{\sqrt{4D_{39}t}} - 1} \right) \times 1000 $$

(6)

The results are shown in Fig. 5. Fig. 5A and 5B show the K isotopic profiles across the diffusion boundary between peridotites and metasomatic melt/fluid at different concentration contrasts (e.g., 10, 50, 100) using either $D_{max}$ or $D_{min}$. As expected, ³⁹K diffuses faster than ⁴¹K from the melt/fluid to the peridotite, which results in lighter K isotope enrichment in the peridotite and heavier K isotope enrichment in the remaining melt/fluid. The higher concentration contrast between the peridotite and metasomatic melt/fluid, the higher degree of K
isotopic fractionation between the two phases (Fig. 5A and 5B). The initial K isotopic compositions of the peridotite and melt/fluid are both equal to that of the BSE, hence there is no need to evoke exotic metasomatic media with abnormal K isotopic compositions (e.g., isotopic light sediments). A locally formed K-enriched melt/fluid with a BSE-like $\delta^{41}$K value can fractionate to yield a broad $\delta^{41}$K range observed in global peridotites.

We further calculated the integrals (from the diffusion boundary to 100 mm in peridotite; the peridotite xenoliths in this study are ~100 mm in size) of the diffusion profiles to represent the K concentrations and isotopic compositions of the modeled “bulk” peridotites (Fig. 5C and 5D). These modeled peridotites have $\delta^{41}$K values varying from -0.77 to -2.75‰ for different concentration contrasts. As expected, the [K] and $\delta^{41}$K values of modeled “bulk” peridotites show a *negative* correlation. The “bulk” peridotites have higher [K] and lower $\delta^{41}$K values after diffusion than the initial compositions. This *negative* correlation between K abundances and isotopes is a diagnostic feature of diffusive K isotopic effects. In contrast, the K isotopic fractionation due to incorporation of a heavy (e.g. due to subduction slab dehydration) fluid would lead to a *positive* correlation between K abundances and $\delta^{41}$K values (see Section 5.2.2).

We note that the integration range (i.e., from the diffusion boundary to 100 mm in peridotite) in this 1D diffusion model are arbitrary. Depending on the “size” of the peridotites analyzed and distance to the diffusion boundaries (i.e., K-rich veins), the large K isotopic fractionation due to diffusion can diminish. This “dilution” effect may be the reason why most of the peridotites are only slightly lighter (e.g., ~-1‰) than the BSE. Extreme values (<-2‰) are rare for bulk peridotites (see the histogram for global non-subduction peridotites in Fig. 5C and 5D). Larger K isotopic fractionation and even lighter K isotopic compositions would be found in smaller rock samples or individual minerals analyzed.

The timescales in these models (Fig. 5 and Supplementary Fig. S1) are arbitrary. We chose a range from 0.1 years to 100,000 years to demonstrate that the light isotopic signatures in peridotite xenoliths can be generated by diffusion in a wide range of parameters (e.g., timescales, temperature-dependent diffusion coefficients, sample sizes, distances to the melt/fluid, and concentration contrasts between the metasomatic melt/fluid and peridotite). However, as we have only obtained bulk K isotope compositions of individual peridotite xenoliths, we are unable to determine the best-fit parameters to constrain the diffusion conditions. In-situ analysis of K isotope profiles across peridotite xenoliths would be necessary to put constraints on this
parameter in natural systems. Nevertheless, the modeling in this study demonstrates that diffusive K isotopic fractionation mechanisms are able to generate the observed lighter K isotope signatures in peridotites.

In summary, similar to other isotope systems, the highly negative δ⁴¹K values in global peridotites, including eight Avacha harzburgites (–1.70 to –0.47‰), are most likely a result of diffusion-driven kinetic isotope fractionation. This conclusion is supported both by laboratory experiments and diffusion modeling discussed above. On the other hand, K mass balance factors make it difficult to explain positive δ⁴¹K that way. Effective K diffusion requires large [K] contrasts between K-rich sources and low-K hosts. Because the K mass in a K-rich source is much greater than the mass of K near the diffusion front, the complementary enrichments in the heavy K isotope in the source are distributed over a much larger K mass, thus small. This is what Ionov and Wang (2021) found in non-subduction zone composite xenoliths from Shava, Tariat: phlogopite-rich veins (former K source) are moderately enriched in heavy K isotopes (–0.75 ± 0.16‰, –0.36 ± 0.09‰, –0.16 ± 0.15‰, –0.15 ± 0.17‰, 0.57 ± 0.06‰) relative to BSE estimates, while bulk xenoliths exhibit significantly light K isotopic compositions (e.g., –2.77 and –1.38‰). Ionov and Wang (2021) interpreted this contrast as the outcome of diffusive K isotopic fractionation: The phlogopite-rich veins represent complementary enrichment in heavy K isotopes, aligning with our expectations in the diffusion modeling. Nevertheless, the phlogopite-rich veins show no positive δ⁴¹K values as high as the subduction zone Avacha harzburgites (up to +1.16‰), which require another mechanism (see **Section 5.2.2**).

Although the diffusion model above demonstrates that the observed lighter K isotopic enrichment in some Avacha peridotites can be attributed to diffusive K isotopic fractionation, we want to note, alternatively, it is possible that the incorporation of isotopically light K sediments could also produce negative δ⁴¹K values in mantle-derived samples (e.g., Hu et al., 2021a). In order to distinguish these two possible explanations: diffusive K isotopic fractionation versus incorporation of isotopically light K sediments, we have conducted a quantitative mixing model following Hu et al. (2021a). As shown in **Fig. 6**, the K isotopic compositions and trace elemental ratios (U/K and Th/Nd) of neither isotopically light nor heavy Avacha peridotites plot on the mixing lines between the mantle and the sediment (or sediment melt). Instead, all the heavy Avacha peridotites plot close to the “slab fluid trend” defined by the K isotopically heavy Martinique lavas from Hu et al. (2021a). These K isotopically heavy Martinique lavas have been
interpreted as the imprint of slab dehydration fluid. Thus we ruled out the influence of subducted sediment components on Avacha peridotites. We will further discuss on the influence of slab dehydration fluid in the following section.

### 5.2.2. Subduction slab dehydration fluid metasomatism

More than half of the Avacha peridotites show positive δ⁴¹K values (+0.22 to +1.16‰; n=11; **Fig. 3**) that have not been found in peridotites from any other locality (Ionov and Wang, 2021). As discussed in *Section 5.2.1*, diffusion-driven kinetic fractionation tends to yield lighter K isotopic values in peridotites, i.e. the opposite to what is observed in these xenoliths. Hence, we link the heavy K isotope enrichments in the Avacha peridotites to their tectonic setting in a subduction zone, i.e. to metasomatic overprint from isotopically heavy fluids associated with subducted slab dehydration.

The Avacha peridotites appear to show two trends on a K isotope vs. [K] plot (**Fig. 4A**), implying that two mechanisms may control the K isotopic variations. The two trends are also seen on a plot of δ⁴¹K vs. K/Nb ratio (**Fig. 4B**). As K is fluid-mobile while Nb is fluid-immobile, the K/Nb ratio can be used as a tracer of fluid activity. While the *negative* trend is most likely due to diffusive isotopic fractionation as discussed above, the *positive* δ⁴¹K vs. K/Nb trend can be best explained by the incorporation of heavy K isotopic signatures from metasomatic fluids. Liu et al., (2020) modeled AOC dehydration in subducted slabs and obtained a δ⁴¹K range of +0.23 to +1.04‰ for the AOC dehydration fluids, which is very close to the δ⁴¹K range for Avacha peridotites with the highest K/Nb (+0.22 to +1.16‰; **Fig. 4B**). In addition, the δ⁴¹K is positively correlated with modal amphibole (**Fig. 4C**), the dominant K carrier in Avacha peridotites (Ionov, 2010), though these modal estimates have high uncertainty. The late-stage, interstitial amphibole precipitated from water-rich melts or fluids expelled from the subducting slab that infiltrated the lithosphere (Bénard et al., 2018; Ionov, 2010) and inherited their heavy K isotopic signature (+0.23 to +1.04‰; Liu et al., 2020; **Fig. 4C**). The low-amphibole Avacha samples generally have lower δ⁴¹K values indicative of diffusion-driven fractionation.

We thus posit that K was added to residual Avacha peridotites by metasomatic media with a range of δ⁴¹K including highly positive values circulating along grain boundaries or in cracks. Metasomatic amphibole-bearing veins ranging from a few cm to micro-cracks (≤1 mm) were found in some Avacha xenoliths (Bénard and Ionov, 2013). Textural data and water content in
minerals for Avacha peridotites (Soustelle et al., 2010) indicate fluid percolation by porous flow and in micro-fractures, which may explain the presence of both low and high δ<sup>41</sup>K values.

These inferences are consistent with previous studies conducted directly on melt inclusions in Avacha peridotites. Ionov et al. (2011) interpreted the enrichment of fluid-mobile trace elements, such as LREE, Sr, Rb, and Ba, as evidence of the contribution from fluids produced by the dehydration of slab materials. Bénard et al. (2018) further narrowed down the source of the fluids to the dehydration products of serpentinized slab rocks, based on U/Th, Pb/Ce, Sr/Nd, and δ<sup>34</sup>S.

The K isotopes of Avacha peridotites are also consistent with published Li-isotope data. Halama et al., (2009) and Pogge von Strandmann et al. (2011) attributed Li isotope variations in Avacha peridotites and veined xenoliths to metasomatism by slab-derived fluids enriched in isotopically heavy Li followed by local diffusive re-equilibration. By contrast isotope compositions of O, Fe and Mg in the Avacha peridotites (Bindeman et al., 2022; Hu et al., 2020a; Weyer and Ionov, 2007) show no effects of metasomatism because of high content of these major elements in the peridotites and low concentrations in slab-related fluids. Thus, K isotopes are a sensitive tracer of slab-derived fluids in melt-depleted peridotites.

**5.3. Potassium isotopes as a tracer for subduction slab components in the mantle wedge**

High-precision K isotope analysis has only become possible within the last few years (Li et al., 2016; Wang and Jacobsen, 2016), yet shortly after its development it was proposed as a potential new tracer for identifying various subducted components in mantle-derived magmas (Parendo et al., 2017). As elaborated in the Introduction, this premise relies on the AOC being enriched in heavier K isotopes compared to the BSE (due to alteration by high-δ<sup>41</sup>K-seawater), and the marine sediments being enriched in lighter K isotopes (due to preferential incorporation of lighter K isotopes into authigenic clays). While this may be partly true, more recent studies reveal that δ<sup>41</sup>K in marine sediments and AOC largely overlap (Fig. 7) (Hu et al., 2020b; Liu et al., 2021; Santiago Ramos et al., 2020). Therefore, contrary to some expectations, these two slab components do not seem to be quantitatively resolvable, either from each other or from the mantle, making it challenging to evaluate their relative contributions using the K isotope tracer.
In addition, subduction processes (e.g., dehydration and metamorphism) can further
fractionate K isotopes in subducting slabs relative to their source compositions. Wang et al.
(2022) showed that subducting marine sediments retain their original K isotopic compositions to
at least a pressure of 2.9 GPa (~90 km depth). Thus, sediment melts released from the slabs into
the mantle wedge preserve the original K isotopic compositions of marine sediments (–1.31 to
–0.02‰; Hu et al., 2020). In contrast, Liu et al. (2020) proposed that the K-bearing hydrous
phases (e.g., phengite) in subducting AOC break down progressively and continually release K-
rich fluids into the overlaying mantle. These dehydration fluids have been modeled to be
enriched in heavy K isotopes (+0.23 ~ +1.04‰; Liu et al. 2020). As such, the K isotope
signatures of sediment melts and AOC dehydration fluids are likely resolvable (–1.31 to –0.02‰
vs. +0.23 to +1.04‰; see schematic illustration in Fig. 8).

Differences in K isotope signatures of sediment melts and AOC dehydration fluids were searched for previously in arc magmas and other mantle-derived lavas (Hu et al., 2021a; Parendo et al., 2022; Sun et al., 2020). Hu et al. (2021a) identified two K isotope trends in the Lesser Antilles arc, and linked lighter lava compositions to additions of up to 5 wt.% sediment melts, while heavier lava compositions were explained by additions of slab dehydration fluids. Parendo et al. (2022) found heavy K isotope enrichment also in Izu arc lavas, which they similarly attributed to slab dehydration fluid. They also observed that the arc front lavas have heavier K isotopes than rear arc lavas, consistent with progressive depletion of the slab in heavy K.

In this study, we found the Avacha arc lava (i.e., peridotite host rocks) show mantle-like K isotopic compositions (–0.53 ± 0.09‰ to –0.34 ± 0.09‰). As shown in Fig. 6 (and supplementary Fig. S2), trace element ratios (e.g., U/K and Th/Nd) and K isotopes of Avacha arc lava plot right on top of the mantle value. The modern Avacha arc lava show no significant influences from either dehydration fluids or sediment melts. Arc lavas are formed during fluid-induced melting in the mantle wedge asthenosphere, so most of K is scavenged from the ambient mantle and enriched in the melt, thus significantly diluting the heavy-K-isotope signatures of dehydration fluids. This results in the mantle-like K isotopic compositions of Avacha arc lava as well as the overall K isotopic compositions of global arc lava being centered around the BSE value (see in Fig. 7).

In contrast, the refractory Avacha peridotites are much more sensitive to this K isotope
signature than arc lavas since the lithospheric mantle is typically K-poor. Indeed, more than half
of the Avacha peridotites analyzed here record this heavy-K-isotope signature, in comparison to
only a small fraction of arc lavas (e.g., Izu arc; **Fig. 7**). The Avacha peridotites undergo
metasomatism due to fluids that have traveled through tens of kilometers of the mantle wedge,
including the base of the lithosphere, while bypassing the melting zone. This is primarily because
these fluids are released at shallower depths compared to the depths necessary for initiating
melting. This interference is also consistent with findings of Parendo et al. (2022) that the arc
front lava closer to the trench have heavier K isotope signatures than rear arc lava (i.e.,
earlier/shallower dehydration fluids are heavier than later/deeper ones).

In summary, for the first time, we observed that sub-arc mantle peridotites have recorded the
dehydration fluid signature previously rarely seen in the arc lava. Our study on Avacha
peridotites show that K isotopes could be a very sensitive tracer of slab processes and subduction
slab dehydration fluids in the mantle. Our study also provides direct evidence that (a) subduction
dehydration and fluid percolation in the mantle lithosphere can significantly fractionate K
isotopes, and (b) slab-derived fluids infiltrated the sub-arc mantle wedge.

**6. Conclusions**

This study provides the first comprehensive elemental and K-isotope data for lithospheric
mantle above a subduction zone based on spinel harzburgite xenoliths from Avacha, Kamchatka.

1) The K isotopic composition of the global lithospheric mantle is highly heterogenous
relative to BSE estimates (−0.43 ± 0.17‰), with δ<sup>41</sup>K ranging from −2.15 to +1.16‰.

2) Mantle peridotites from various tectonic settings are commonly enriched in lighter K isotopes compared to the BSE, most likely due to diffusion-driven kinetic isotopic fractionation during metasomatism by K-rich media.

3) By contrast, subduction zone peridotites from Avacha and possibly elsewhere typically show positive δ<sup>41</sup>K (up to +1.16 ± 0.08‰), which is best explained by incorporation of isotopically heavy dehydration fluids expelled from subducting slabs. This is consistent with K isotope data on arc lavas and eclogites as well as modeling of isotopic fractionation on dehydration.

4) This study provides direct evidence that slab-derived fluids infiltrate arc mantle. Isotopes of potassium, a fluid-mobile trace element, are sensitive to metasomatic activity in arc mantle and may be used as a novel tracer of dehydration fluids from subducted slabs.
Acknowledgements

KW acknowledges financial support from the McDonnell Center for the Space Sciences and NASA (Emerging Worlds Program grant number #80NSSC21K0379). We thank P. Plechov for providing Sample Av-50. K.W. thanks Michael J. Krawczynski for helpful discussion. Reviews by two anonymous referees and editorial handling by Dr. Hickey-Vargas are gratefully acknowledged.
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<table><caption>Table 1 Potassium isotope compositions of peridotite xenoliths and host rocks from Avacha Volcano, Kamchatka</caption><thead><tr><th></th><th>Mass (mg)</th><th>Sample type</th><th>[K] (ppm)</th><th>δ<sup>41</sup>K<sub>NIST3141a</sub></th><th>95% C.I.<sup>a</sup></th><th>n<sup>b</sup></th></tr></thead><tbody><tr><td colspan="7"><strong>Peridotite xenoliths</strong></td></tr><tr><td>Av-1</td><td>293.4</td><td>Harzburgite, whole-rock</td><td>8</td><td>0.24 ± 0.02</td><td></td><td>12</td></tr><tr><td>Av-2</td><td>182.1</td><td>Harzburgite, whole-rock</td><td>14</td><td>1.16 ± 0.08</td><td></td><td>6</td></tr><tr><td>Av-5</td><td>281.3</td><td>Harzburgite, whole-rock</td><td>12</td><td>0.40 ± 0.02</td><td></td><td>12</td></tr><tr><td>Av-6</td><td>186.6</td><td>Harzburgite, whole-rock</td><td>19</td><td>0.63 ± 0.06</td><td></td><td>10</td></tr><tr><td>Av-6 (literature)<sup>c</sup></td><td>253.5</td><td>Harzburgite, whole-rock</td><td>17</td><td>0.62 ± 0.04</td><td></td><td>11</td></tr><tr><td>Av-8</td><td>172.2</td><td>Harzburgite, whole-rock</td><td>27</td><td>0.24 ± 0.08</td><td></td><td>10</td></tr><tr><td>Av-8 (literature)<sup>c</sup></td><td>260.0</td><td>Harzburgite, whole-rock</td><td>21</td><td>0.40 ± 0.05</td><td></td><td>11</td></tr><tr><td>Av-9</td><td>168.6</td><td>Harzburgite, whole-rock</td><td>9</td><td>0.22 ± 0.16</td><td></td><td>5</td></tr><tr><td>Av-10</td><td>120.2</td><td>Harzburgite, whole-rock</td><td>13</td><td>-1.70 ± 0.19</td><td></td><td>8</td></tr><tr><td>Av-12</td><td>151.3</td><td>Harzburgite, whole-rock</td><td>8</td><td>-0.49 ± 0.07</td><td></td><td>7</td></tr><tr><td>Av-13</td><td>184.5</td><td>Harzburgite, whole-rock</td><td>13</td><td>0.77 ± 0.04</td><td></td><td>12</td></tr><tr><td>Av-13 (replicate)<sup>d</sup></td><td>165.0</td><td>Harzburgite, whole-rock</td><td>17</td><td>0.49 ± 0.04</td><td></td><td>12</td></tr><tr><td>Av-14</td><td>133.2</td><td>Harzburgite, whole-rock</td><td>17</td><td>0.49 ± 0.11</td><td></td><td>10</td></tr><tr><td>Av-15</td><td>143.9</td><td>Harzburgite, whole-rock</td><td>6</td><td>-0.53 ± 0.06</td><td></td><td>7</td></tr><tr><td>Av-33</td><td>112.6</td><td>Harzburgite, whole-rock</td><td>27</td><td>-1.20 ± 0.13</td><td></td><td>7</td></tr><tr><td>Av-33 (replicate)<sup>d</sup></td><td>215.4</td><td>Harzburgite, whole-rock</td><td>32</td><td>-0.98 ± 0.04</td><td></td><td>11</td></tr><tr><td>Av-41</td><td>133.0</td><td>Harzburgite, whole-rock</td><td>19</td><td>-1.15 ± 0.06</td><td></td><td>12</td></tr><tr><td>Av-G-1</td><td>134.4</td><td>Harzburgite, whole-rock</td><td>9</td><td>-0.70 ± 0.06</td><td></td><td>7</td></tr><tr><td>Av-G-2</td><td>123.3</td><td>Harzburgite, whole-rock</td><td>8</td><td>-0.47 ± 0.09</td><td></td><td>8</td></tr><tr><td colspan="7"><strong>Host magma products</strong></td></tr><tr><td>Av-18 ground</td><td>3.7</td><td>Bulk selvage<sup>e</sup>, ground</td><td>1660</td><td>-0.53 ± 0.09</td><td></td><td>11</td></tr><tr><td>Av-18 crystal</td><td>5.8</td><td>Amphibole in selvage<sup>e</sup></td><td>1722</td><td>-0.51 ± 0.09</td><td></td><td>22</td></tr><tr><td>Av-50 selvage</td><td>6.6</td><td>Amphibole selvage<sup>e</sup></td><td>1521</td><td>-0.44 ± 0.07</td><td></td><td>21</td></tr><tr><td>Av-50 amphibole</td><td>7.2</td><td>Amphibole in magmatic rock<sup>f</sup></td><td>1607</td><td>-0.50 ± 0.11</td><td></td><td>17</td></tr><tr><td>Av-39 amphibole</td><td>17.6</td><td>Amphibole selvage<sup>e</sup>, crystals</td><td>1992</td><td>-0.34 ± 0.09</td><td></td><td>22</td></tr></tbody></table>

<footnote><sup>a</sup> 95% confidence interval (c.i.): z × [σ/√(n-1)], where z is the student's t number, σ = standard deviation, and n = number of measurements.</footnote>

<footnote><sup>b</sup> Number of measurements.</footnote>
856 c Literature values are from Ionov and Wang (2021) for samples digested in Parr high-pressure vessels (different aliquots of the same whole-rock powders).

857 d Samples in this study dissolved in Parr high-pressure digestion vessels. All other samples in this study were dissolved using the hot-plate method.

858 e Black coatings on xenolith surfaces deposited from host andesitic magma consisting of mm-size euhedral amphibole crystals, silicate glass and minor olivine.

860 f Av-50 is a magmatic cumulate dominated by coarse black amphibole crystals (hand-picked for analysis).

863
<table><caption>Table 2 Potassium isotope compositions of geological reference materials in this study and literatures</caption><thead><tr><th></th><th>Mass (mg)</th><th>Type</th><th>δ<sup>41</sup>K<sub>NIST3141a</sub></th><th>95% C.I.<sup>a</sup></th><th>n<sup>b</sup></th></tr></thead><tbody><tr><td colspan="6"><strong>This study</strong></td></tr><tr><td>BHVO-2 #1 Day 1</td><td></td><td>basalt</td><td>USGS -0.56 ± 0.06</td><td></td><td>9</td></tr><tr><td>BHVO-2 #1 Day 2</td><td></td><td>basalt</td><td>USGS -0.55 ± 0.08</td><td></td><td>9</td></tr><tr><td>BHVO-2 #1 Day 3</td><td></td><td>basalt</td><td>USGS -0.54 ± 0.16</td><td></td><td>11</td></tr><tr><td>BHVO-2 #1 Day 4</td><td></td><td>basalt</td><td>USGS -0.57 ± 0.10</td><td></td><td>12</td></tr><tr><td>BHVO-2 #1 Day 5</td><td></td><td>basalt</td><td>USGS -0.59 ± 0.14</td><td></td><td>12</td></tr><tr><td>BHVO-2 #1 Day 6</td><td></td><td>basalt</td><td>USGS -0.36 ± 0.05</td><td></td><td>10</td></tr><tr><td>BHVO-2 #1 Day 7</td><td></td><td>basalt</td><td>USGS -0.44 ± 0.11</td><td></td><td>11</td></tr><tr><td>BHVO-2 #1 Day 8</td><td></td><td>basalt</td><td>USGS -0.47 ± 0.08</td><td></td><td>8</td></tr><tr><td>BHVO-2 #1 Day 9</td><td></td><td>basalt</td><td>USGS -0.51 ± 0.03</td><td></td><td>12</td></tr><tr><td>BHVO-2 #1 Day 10</td><td></td><td>basalt</td><td>USGS -0.48 ± 0.10</td><td></td><td>11</td></tr><tr><td>BHVO-2 #1 Day 11</td><td></td><td>basalt</td><td>USGS -0.49 ± 0.10</td><td></td><td>4</td></tr><tr><td>BHVO-2 #1 Day 12</td><td></td><td>basalt</td><td>USGS -0.46 ± 0.06</td><td></td><td>10</td></tr><tr><td>BHVO-2 #1 Day 13</td><td></td><td>basalt</td><td>USGS -0.48 ± 0.09</td><td></td><td>11</td></tr><tr><td>BHVO-2 #1 Day 14</td><td></td><td>basalt</td><td>USGS -0.52 ± 0.07</td><td></td><td>8</td></tr><tr><td>BHVO-2 #2<sup>c</sup></td><td></td><td>basalt</td><td>USGS -0.44 ± 0.15</td><td></td><td>9</td></tr><tr><td>BHVO-2 #3<sup>c</sup></td><td></td><td>basalt</td><td>USGS -0.44 ± 0.07</td><td></td><td>7</td></tr><tr><td>BHVO-2 #4<sup>c</sup></td><td>~0.01</td><td>basalt</td><td>USGS -0.43 ± 0.06</td><td></td><td>7</td></tr><tr><td>BHVO-2 #5<sup>c</sup></td><td>~0.1</td><td>basalt</td><td>USGS -0.48 ± 0.09</td><td></td><td>9</td></tr><tr><td>G-2 #1<sup>c</sup></td><td>9.7</td><td>granite</td><td>USGS -0.53 ± 0.15</td><td></td><td>9</td></tr><tr><td colspan="6"><strong>Literatures<sup>d</sup></strong></td></tr><tr><td>BHVO-2</td><td></td><td>basalt</td><td>USGS -0.46 ± 0.09</td><td></td><td>108</td></tr><tr><td>G-2</td><td></td><td>granite</td><td>USGS -0.45 ± 0.04</td><td></td><td>79</td></tr></tbody></table>

<footnote>a 95% confidence interval (c.i.): z × [σ/√(n-1)], where z is the student's t number, σ = standard deviation, and n = number of measurements.</footnote>

<footnote>b Number of measurements.</footnote>

<footnote>c These geological reference materials were dissolved, chemically processed, and analyzed along with mantle samples as unknowns.</footnote>

<footnote>d Recommended average in Wang et al. (2021) compiled from literature values (see Wang et al., 2021 for references).</footnote>
<figure><img src="image_2.png" /><figcaption>**Figure 1** The location of Avacha volcano in Kamchatka arc. Other active volcanos are plotted as triangles. The volcano data are from ASTER Volcano Archive – NASA-META. Figure made with GeoMapApp (www.geomapapp.org) / CC BY / CC BY (Ryan et al., 2009).</figcaption></figure>
<figure><img src="image_3.png" /><figcaption>**Figure 2.** Photograph of whole-rock specimen (A) and thin-section photomicrographs (B and C) of peridotite xenolith from Avacha (Ionov, 2010). The outer parts of the xenolith (several cm away from contacts with host andesite) were removed, and only the core material shown here was used for the whole-rock powder. The core material contains no alteration products or host magma veins. The major minerals olivine (ol) and orthopyroxene (opx), as well as accessory minerals spinel (sp) and trace interstitial amphibole (am) are labeled.</figcaption></figure>

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<figure><img src="image_4.png" /></figure>

**Figure 3** The K isotopic compositions of subduction zone (Avacha) and global non-subduction zone peridotites and crustal rocks. The non-subduction zone peridotites are from the Siberian craton, Central Asian Orogenic Belt, Eifel volcanic field in Germany and the Gnotuk maar in southeastern Australia (Ionov and Wang, 2021). The subduction zone peridotites in this study are all from Avacha Volcano, Kamchatka arc. The MORB (mid-ocean ridge basalt) data and the BSE (Bulk Silicate Earth) value are from (Tuller-Ross et al., 2019a). All samples in this figure including literature data were analyzed in the same laboratory using the same method.

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<figure><img src="image_5.png" /></figure>

**Figure 4** (A) K isotopic compositions versus K concentrations in Avacha peridotites. All isotopic and K concentration data are from this study. (B) K isotopic compositions versus K/Nb ratios in Avacha peridotites. The Nb concentrations are from Ionov (2010); not all samples in this study have Nb data available. The K isotopic range of modelled dehydration fluids removed from subducting altered oceanic crust, as well as diffusion isotopic trend are shown for comparison (Liu et al., 2020). (C) K isotopic compositions versus amphibole modal abundances (Ionov, 2010) in Avacha peridotites. Note that amphibole abundances are not available for all samples in this study.

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<figure><img src="image_6.png" /></figure>

**Figure 5** (A) D<sub>max</sub> Scenario: Diffusive K isotopic fractionation profiles from a metasomatic melt/fluid to peridotite at different concentration contrasts. (B) D<sub>min</sub> Scenario: Diffusive K isotopic fractionation profiles from a metasomatic melt/fluid to peridotite at different concentration contrasts. (C) D<sub>max</sub> Scenario: The K concentrations and isotopic compositions of modeled “bulk” peridotites are integrated from -100 mm to the boundary. (D) D<sub>min</sub> Scenario: The K concentrations and isotopic compositions of modeled “bulk” peridotites are integrated from -100 mm to the boundary. As illustrated by the arrow, diffusive isotopic fractionation will create a
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negative trend between bulk K concentrations and isotopic compositions. The histogram on the y-axis shows the variation of K isotopic compositions of 35 whole-rock analyses of non-subduction zone peridotites from Ionov and Wang (2021). We interpret those negative values as the result of diffusive kinetic fractionation. Note that the temperature-dependent diffusion coefficient, *D*, and time scale, *t*, are trade-offs. With different combinations of reasonable diffusion coefficients and timescales, we can account for the observed variations in light K isotope compositions found in Avacha peridotites.
<figure><img src="image_7.png" /></figure>

**Figure 6** K isotopic compositions versus 10000U/K (A) and Th/Nd (B) in Avacha arc lava (i.e., host rocks) and peridotites. The slab fluid trend is from Hu et al. (2021a), where the slope is defined by K isotopically heavy Martinique lavas (interpreted as fluid input; Hu et al., 2021a). The binary mixing lines between mantle and Kamchatka sediments (or sediment melts) are calculated in this study. The U, K, Th, and Nd concentrations of the mantle are from Palme and O'Neill (2014). The mantle $\delta^{41}$K value is from Tuller-Ross et al. (2019a). The U, K, Th, and Nd concentrations of the Kamchatka sediments are from Plank (2014). The $\delta^{41}$K values of Kamchatka sediments are not known. We use the lowest $\delta^{41}$K value (–1.31‰) among global sediments from Hu et al., (2020b) in order to compare to the K isotopically light peridotites. The U, K, Th, and Nd concentrations of bulk Avacha arc lava are from Ishimaru et al. (2007) and their $\delta^{41}$K values are the average $\delta^{41}$K of host rock samples in this study. The measured Th/Nd ratios of the bulk Avacha arc lava are corrected for the effect of 15% partial melting, using the partition coefficients from Niu and Hékinian (1997) and from https://earthref.org/. Our calculation follows Hu et al. (2021a) and Labanieh et al. (2012). The Kamchatka sediment melt Th/Nd ratio was calculated by 50% partial melting of the Kamchatka sediments using the partition coefficients from Johnson and Plank (2000) at 800°C.

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<figure><img src="image_8.png" /></figure>

**Figure 7** Histograms of K isotopic compositions of altered oceanic crust, marine sediments, arc lavas from the literature, and mantle wedge (Avacha peridotites) from this study. The K isotopic range for modelled dehydration fluids removed from subducting altered oceanic crust are shown as a blue-grey field (Liu et al., 2020). Literature data: altered oceanic crust (Liu et al., 2021; Parendo et al., 2017; Santiago Ramos et al., 2020), marine sediment (Hu et al., 2020b), arc lava (Hu et al., 2021a; Parendo et al., 2022).
<figure><img src="image_9.png" /><figcaption>**Figure 8** Cartoon illustrating the subduction process and associated K isotopic fractionation.</figcaption></figure>

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