Extremely light K in subducted low-T altered oceanic crust: Implications for K recycling in subduction zone

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Extremely light K in subducted low-T altered oceanic crust: Implications for K recycling in subduction zone Haiyang Liu a,b,c,d, Kun Wang(王昆) d,⇑, Wei-Dong Sun a,b,c,e,⇑, Yilin Xiao f, Ying-Yu Xue a,b,c, Brenna Tuller-Ross d b

a Center of Deep Sea Research, Institute of Oceanology, Chinese Academy of Sciences, Qingdao 266071, PR China Laboratory for Marine Mineral Resources, Qingdao National Laboratory for Marine Science and Technology, Qingdao 266237, PR China c Center for Ocean Mega-Science, Chinese Academy of Sciences, Qingdao 266071, PR China d Department of Earth and Planetary Sciences, McDonnell Center for the Space Sciences, Washington University in St. Louis, St. Louis, MO 63130, USA e University of Chinese Academy of Sciences, Beijing 100049, PR China f CAS Key Laboratory of Crust-Mantle Materials and Environments, CAS Center for Excellence in Comparative Planetology, School of Earth and Space Sciences, University of Science and Technology of China, Hefei 230026, PR China

Received 4 June 2019; accepted in revised form 20 March 2020; Available online 4 April 2020

Abstract In order to investigate the behavior of potassium (K) isotopes during metamorphic dehydration and to further constrain its implications for K recycling in the subduction zone, we measured K isotopic compositions of whole-rocks and separated minerals (phengite, omphacite, amphibole) from Sumdo eclogites, Tibet, using a recently established high-precision analysis method. Our data reveal that the d41K (‰) of the whole-rock eclogites (
1.64 to
0.24) displays dramatically lower values than observed in fresh mid-ocean ridge basalts (MORB) (
0.43 ± 0.17) and altered MORB (
0.76 to
0.11). In addition, the d41K values of eclogites show a positive correlation with both K2O contents and K/Nb ratios, which suggests that the low d41K values were most likely caused by dehydration during subduction. Thus, isotopically heavy K may be released into the mantle wedge, while the light component is subducted into the deep mantle. Therefore, the K isotope systems have the potential to trace subducted crustal materials and to create heterogeneity within the mantle. Mineral separates from Sumdo eclogites are highly heterogeneous in K isotopic compositions, ranging from
0.98 to +0.23 in phengite, from
1.25 to
1.10 in omphacite and from
1.16 to
0.09 in amphibole. The K isotope fractionations between amphibole and phengite and between amphibole and omphacite vary from
0.30 to +0.25 and from
0.04 to +0.63, respectively, indicating K isotopic disequilibrium between amphibole and omphacite/phengite, which might result from the multistage growth of minerals during subduction metamorphism. However, generally phengites show heavier K isotopic compositions than the coexisting omphacite (D41Kphengite-amphibole = +0.25) and amphibole (D41Kphengite-amphibole = +0.03  +0.30, except TB193 is
0.25), which may imply that the K isotopic fractionation is controlled by the difference in coordination numbers of K between phengite (6) and omphacite (7 to 8)/amphibole (8). Ó 2020 Elsevier Ltd. All rights reserved. Keywords: K isotope; Fractionation; Subduction dehydration; Eclogite; Tibet

⇑ Corresponding authors at: Center of Deep Sea Research, Institute of Oceanology, Chinese Academy of Sciences, Qingdao 266071, PR

China (W.-D. Sun). E-mail addresses: [email protected] (K. Wang(王昆)), [email protected] (W.-D. Sun). https://doi.org/10.1016/j.gca.2020.03.025 0016-7037/Ó 2020 Elsevier Ltd. All rights reserved.

H. Liu et al. / Geochimica et Cosmochimica Acta 277 (2020) 206–223

1. INTRODUCTION Potassium (K) is the eighth most abundant element in the Earth’s crust (K2O = 1.81 wt%), fifteenth in the bulk Earth (K = 550 mg/g), sixth in seawater (K = 399 mg/g) and is also a critical nutrient in biological activities, making it a key element to track a variety of processes (McDonough and Sun, 1995; Riley and Tongudai, 1967; Rudnick and Gao, 2014; Sardans and Pen˜uelas, 2015). As an incompatible large ion lithophile element, K tends to partition into melt during mantle melting (Arevalo et al., 2009; McDonough and Sun, 1995; McDonough et al., 1992). Moreover, K has two stable isotopes, 39K (93.2581%) and 41K (6.7302%), and one radioactive isotope 40 K (0.0117%) which has a half-life of 1.277  109 years. While it has been historically difficult to measure K isotopes, the recently developed high-precision analysis method has made it possible to resolve the behavior of K isotopic variations (41K/39K < 0.1‰) during various geological processes (Chen et al., 2019; Hu et al., 2018; Li et al., 2020, 2019a, 2016b; Morgan et al., 2018; Santiago Ramos et al., 2018; Sun et al., 2019; Teng et al., 2020; Wang and Jacobsen, 2016a, b; Xu et al., 2019). Potassium is among the most sensitive indicators of water-rock exchange during seafloor hydrothermal alteration (Hart and Staudigel, 1982; Kelley et al., 2003; Staudigel, 2014; Staudigel et al., 1996; Zhang and SmithDuque, 2014). Moreover, the large K isotopic composition offset between Bulk Silicate Earth and seawater (0.6‰ in d41K values) implies that the cycling of K among different geological reservoirs is associated with significant K isotope fractionation (Li et al., 2020, 2019a, 2016b; Morgan et al., 2018; Santiago Ramos et al., 2018; Teng et al., 2020; Wang and Jacobsen, 2016a). For instance, studies on in situ oceanic crust and hydrothermal vent fluids have demonstrated that K is leached from the basalts during hightemperature reactions, while taken up by basalts during low-temperature alterations, which was also confirmed by subsequent experimental work (Hart and Staudigel, 1982; Kelley et al., 2003; Seyfried and Bischoff, 1979; Staudigel, 2014; Staudigel et al., 1996; Zhang and Smith-Duque, 2014). In addition, recent investigation has revealed that ophiolites from the Bay of Islands (i.e., fragments of ancient oceanic crust) display heavier K isotopic compositions (up to 0.67‰ in d41K values) than the fresh oceanic crust and d41K values are well correlated with 87Sr/86Sr, which can be explained by K isotope addition during hydrothermal alteration (Parendo et al., 2017). A direct investigation on altered oceanic crust (AOC) samples drilled at ODP Site 801 (ca. 170 Ma) also documented a large variation in d41K, which may reflect the uptake of isotopically heavy K (d41K  0.1‰, relative to NIST SRM3141a) from seawater. Therefore, the low-T AOC entering the subduction zone contains generally higher K concentrations and has heavier K isotopic compositions than the fresh oceanic crust and the normal mantle (d41K 
0.43 ± 0.17‰, 2sd, Tuller-Ross et al., 2019). Based on these lines of evidence, Parendo et al. (2017) suggest that K isotopic systematics can be used as an effective tracer of oceanic crust subducted into the deep mantle.

207

Notably, the investigation of K isotopic characteristics of the continental basaltic lavas from Northeast China shows evidence that K isotopes have the potential to trace different types of recycled crustal materials in the mantle (Sun et al., 2019). However, K is also fluid mobile in subduction zones and strong depletions of K (>40%, also Rb and Ba) have been demonstrated by the trace element compositions of the eclogites, blueschists and mafic granulites from numerous high-pressure terranes when compared with their protoliths (e.g., unaltered and altered MORB or ocean island basalts, Becker et al., 2000; Schmidt and Poli, 2014; Schmidt, 1996; Zack et al., 2001). Hence, large K isotopic fractionation similar to that of lithium and boron isotopes (Marschall et al., 2007; Peacock and Hervig, 1999) is expected to occur during metamorphic dehydration in a subduction zone, which, however, has not yet been determined. In addition, K isotopes could also be fractionated between minerals in eclogites because the coordination number for K is different among major K-rich minerals in eclogite, with the coordination number of K being 6 in phengite (Li et al., 2019b), 7–8 in omphacite (Hawthorne and Calvo, 1977; Huebner and Papike, 1970) and 8 in amphibole (Huebner and Papike, 1970). Theory predicts that at equilibrium, heavier isotopes (e.g., 7Li, 26Mg and 56Fe) will prefer high forceconstant bonds that are generally associated with lower coordination numbers (Schauble, 2004; Li et al., 2016a, 2011b, 2019a; Sun et al., 2016). However, the intermineral K isotopic fractionation remains to be determined, which is a prerequisite to apply K isotopes as tracers for geological processes. In this regard, we carried out highprecision K isotopic analyses on whole rock and mineral separates from Sumdo eclogites, Tibet, which were derived from altered MORB and experienced metamorphic dehydration during subduction (Liu et al., 2019a; Liu et al., 2019b). These samples thus provide an excellent opportunity to investigate the K isotope systematics of a deeply subducted low-T AOC and may ultimately improve our ability to use the K isotope system to track recycled crustal materials in the mantle. 2. GEOLOGICAL BACKGROUND AND SAMPLES The Tibetan Plateau is the youngest and classic subduction related, continent-continent collisional orogen that was largely created by the Indo-Asian collision in a succession of spreading, subduction, and collision of juvenile oceanic material and variably sized continental fragments over the past 70–50 Ma (Sun et al., 2018a, 2018b; Zhang et al., 2013). From south to north, the interior of the Tibetan orogenic belt was divided into the Himalaya, Lhasa, Qiangtang, and Songpan-Ganzi terranes (Li et al., 2015; Meng et al., 2019; Sun et al., 2018a, 2018b; Zhang et al., 2013). The Lhasa terrane, located in the southern Tibet, is bounded by the Yarlung-Zangbo suture to the south and the Bangong-Nujiang suture to the north (Pan et al., 2006; Sun et al., 2018c; Yang et al., 2009; Zhu et al., 2011, 2010). The recently recognized Sumdo high-pressure metamorphic belt (e.g., eclogite, blueschist) located in the central

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Lhasa terrane, together with the Paleo-Tethys ophiolite, Carboniferous to Permian island arc volcanic rocks, Permian orogenic granite and the regional angular unconformity between the Middle and Upper Permian, reveal the existence of a subduction zone, which was resulted from the northward subduction of the Paleo-Tethys beneath the North Lhasa terrane (Cao et al., 2017b; Cheng et al., 2015; Liu et al., 2019b, 2016; Wang et al., 2019; Weller et al., 2016; Yang et al., 2009; Zeng et al., 2009). Previous researches have demonstrated that the Sumdo eclogite was derived from an N-MORB (Li et al., 2009) or ocean island basalt (OIB) (Cheng et al., 2015) source and formed in a back-arc setting (Cheng et al., 2015; Yang et al., 2009; Zeng et al., 2009). Based on the trace element signatures and positive correlation between initial 87Sr/86Sr and d18O values, Liu et al. (2019a) and Zeng et al. (2009) suggest that the protolith of the Sumdo eclogite had experienced low temperature seawater alteration prior to subduction. The whole-rock Sm-Nd dating (306 ± 50 Ma, Li et al., 2009) and in situ zircon dating (304 ± 5 Ma, Cheng et al., 2015) indicated that the protoliths were formed at 305 Ma. The peak P-T conditions of the Sumdo eclogite are estimated at 2.6–3.3 GPa and 630–780 °C (Li et al., 2007; Liu et al., 2019b; Weller et al., 2016; Yang et al., 2009; Zeng et al., 2009) or 2.5 GPa and 520 °C (Cao et al., 2017a). Evidenced by the zircon in situ dating, the eclogite-facies metamorphic age has been constrained at 260 Ma (Cao et al., 2017a; Weller et al., 2016; Yang et al., 2009). Metamorphic P-T estimation and mineral chemistry data suggested that after the peak metamorphism, the Sumdo eclogite experienced a heating stage during early exhumation (Cao et al., 2017a, 2017b; Liu et al., 2019b; Weller et al., 2016). The eclogite samples investigated in this work were collected from the Sumdo-Gongboyamda region and have previously been the subjects of comprehensive studies of petrography, mineral chemistry, fluid inclusions, radiogenic (Sr-Nd-Pb) isotopes and stable (oxygen, lithium) isotopes by Liu et al. (2019a) and Liu et al. (2019b). These eclogites are massive and consist mainly of garnet, omphacite, amphibole, rutile, phengite, zoisite and epidote/clinozoisite (Tables S1). Detailed petrographic observations indicate that the dominant minerals formed during prograde and peak metamorphic stages are garnet, omphacite, phengite, rutile, and amphibole (Representative petrographic photomicrographs are shown in Fig. S1, Liu et al., 2019b). In addition, some eclogite samples experienced various degrees of retrograde metamorphism, represented by the retrograde fine-grained amphibole and epidote-group minerals (e.g., epidote, zoisite and clinozoisite, Liu et al., 2019b). The garnets commonly contain a core with plentiful mineral inclusions (e.g., quartz, biotite, and apatite), which were formed in the early prograde stage. The garnets constitute more than 99 mol. % of almandine, pyrope, and grossular, whereas the K concentrations in garnet are far below the detection limit (Liu et al., 2019b). Omphacites occur as both porphyroblasts and inclusions in garnet. The K concentrations of the porphyroblast omphacite are marginally higher than the detection limit, whereas the omphacites included in garnet contain much higher K concentrations

than the porphyroblast omphacites (Liu et al., 2019b). The phengites, observed in the matrix (i.e., minerals other than the mineral inclusions), display high concentration in both SiO2 (48.6–53.7 wt%, corresponding to Si pfu: 3.4– 3.6 per 11 oxygens) and K2O (9.8–11.0 wt%). Therefore, the white micas investigated in these samples can all be identified as phengite (Liu et al., 2019b). The Sumdo eclogites contain several stages of amphibole, the earlier generations of amphiboles occur as inclusions in garnet cores or as porphyroblasts in the matrix in equilibrium with HP minerals, respectively, indicating that they were formed during the prograde and peak metamorphic stage. While, the later generations of amphiboles occur as retrograde reaction rims around garnet and omphacite, suggesting that they are retrograded minerals (Liu et al., 2019b). Different stages of amphibole have a wide compositional range but are all low in K concentrations (<0.20 wt%, Liu et al., 2019b). The chemical compositions of the retrograded amphibole were inherited from the original minerals (e.g., garnet and omphacite). Similarly, several stages of epidote-group minerals were identified in Sumdo eclogites, which all have low K concentrations (below the detection limit). Collectively, the phengite, omphacite and amphibole are the major host for K in the Sumdo eclogite (Liu et al., 2019b). In addition, the direct country rocks of the Sumdo eclogite body are garnet-bearing mica-quartz schists, which are fine- to medium-grained and consist mainly of quartz (60%), muscovite (25%), feldspar (10%) and minor garnet (5%). Previous investigations suggested that the quartz schists were not involved in the high-grade metamorphism, but experienced amphibolites-facies retrograde metamorphism, simultaneously with the Sumdo eclogite (Yang et al., 2009; Li et al., 2011a, 2012). In this study, twenty eclogites, one country rock (garnet-bearing mica-quartz schists) and 13 K-rich minerals (4  phengite, 3 omphacite, 6  amphibole, we acknowledge that the amphibole separates are a mixture of prograde and retrograde amphiboles) from 6 samples were investigated for K isotopic analyses in order to determine the K isotope systematics in subduction zones. 3. ANALYTICAL METHODS 3.1. Whole rock major elements and mineral chemistry Whole rock major elements and Cr, Sr, V and Ni were determined on fused glass discs with an X-Ray Fluorescence Spectrometer (XRF) at the laboratory of ALS minerals at Guangzhou. Pre-ignition was used to obtain the loss on ignition (LOI) prior to major element analyses. The analytical precision and accuracy for the major elements are better than ± 1%. The chemical compositions of the omphacite included in garnet were determined by a laser ablation ICP MS system (Element 2), with an internal standard-independent calibration strategy, at the University of Go¨ttingen, Germany. Each measurement incorporated a 20 s background acquisition followed by 40 s data acquisition from the sample. NIST 610, StHs6/80-G, and GSD-1G were measured every 4 samples in order to correct the time-dependent drift of

H. Liu et al. / Geochimica et Cosmochimica Acta 277 (2020) 206–223

sensitivity and mass discrimination. Detailed analytical procedures were described in Wu et al. (2018). 3.2. Potassium isotopes For the K isotope measurement, approximately 10– 100 mg of the whole rock powders and mineral grains (e.g., phengite, omphacite, amphibole) of each sample were weighed and separated into Teflon beakers, with the exact mass depending on each sample’s K concentration. First, the whole rock samples were dissolved in ultra-pure concentrated HF + HNO3 under heat lamps for at least 5 days in order to completely dissolve the samples. While for the mineral separates, they were digested into high pressure Teflon bombs along with the HF + HNO3 and placed in an oven for at least 7 days at 200 °C to guarantee completely dissolution. Next, the dissolved samples were evaporated and re-digested in ultra-pure 6 mol/L HCl for at least 24 hours under heat lamps. Finally, 0.7 mol/L ultra-pure HNO3 was added to re-dissolve the digested samples before they were ready to be loaded into the ion exchange columns. K was purified from the silicate matrix by a two-step liquid chromatography procedure in a clean lab at Department of Earth and Planetary Sciences, Washington University in St. Louis following the procedure described by Chen et al. (2019) and Wang and Jacobsen (2016a). Both columns were filled with Bio-Rad AG 50W-X8 resin (100– 200 mesh) with different volumes to completely separate the sample matrix ions (e.g., Cr, Na, Rb). The second column procedure was carried out two or three times in the case of high matrix ion/K ratio (which was the case for some of the mineral separates). At least one reference material (e.g., BHVO-2, BCR-2, G-2 and AGV-2) was processed through column chemistry with each batch of unknown samples. All separations were monitored with Thermo Scientific iCAP Q quadrupole ICP-MS analysis to guarantee both high K yield (>99% recovery) and low matrix ion/K ratios (Tables S2 and S3). The K concentrations of the samples measured in this study agree well with the published data within 20% (most within 10%), which also demonstrates the high recovery of K during purification (Tables S2 and S3). The total procedure blank in this study is 13 ng, which is negligible compared to the amount of K (100 mg) loaded onto the columns. K isotopic compositions were measured by the samplestandard (NIST3141a) bracketing method on a Thermo Scientific Neptune Plus MC-ICP MS equipped with an APEX X high sensitivity desolvation system at Department of Earth and Planetary Sciences, Washington University in St. Louis (Tables S4). By optimizing the instrument settings of Neptune Plus MC-ICP MS and APEX X, the interference signals of argon-hydride and 40Ar+ are successfully reduced (Chen et al., 2019). Generally, a mean 39K signal of about 15 V (500 ppb) was routinely obtained. K isotopic composition data are presented in delta notation in per mil (‰) relative to a pure K standard solution NIST SRM3141a that is defined as d41K = [(41K/39K)sample/ (41K/39K)NIST3141a
1]  1000. K isotopic fractionation between minerals are reported by D41KA-B = d41 KA - d41KB, where A and B represent different mineral

209

phases. Based on the long-term repeated analysis on standards, the external precision is better than 0.1‰ (95% confidence interval; 95% CI). The K isotopic compositions of the rock reference materials (e.g., BHVO-2, BCR-2, G-2 and AGV-2) analyzed during the course of this study (Tables S2) agree well within uncertainties with published values (Chen et al., 2019; Li et al., 2016b; Morgan et al., 2018; Wang and Jacobsen, 2016a; Xu et al., 2019), confirming that our data are accurate. 4. RESULTS The whole rock major and trace elements and isotope compositions are shown in Tables 1 and S5, S6. The major element concentrations of 10 eclogite and 1 country rock samples were measured in this study (Table S5). The major element data of other samples and all the trace elements along with Sr-Nd-Pb-O and Li isotope compositions (Tables S5 and S6) are from Liu et al. (2019a) and Liu et al. (2019b). 4.1. Major elements The investigated eclogite samples show a limited range in major elements (e.g., SiO2 = 39.5–49.1 wt%, with an average of 45.7 ± 3.2 wt%, Table S5), corresponding to the typical major element compositions of N-MORB, which is consistent with the previous conclusions (Liu et al., 2019a, 2019b; Zeng et al., 2009). They display low concentrations of K (415–9132 mg/g, with an average of 1756 ± 3842 mg/g, Fig. 1) and total alkalis (K2O + Na2O = 0.61–2.73 wt%, with an average of 1.78 ± 1.17 wt%, Table S5), with Na2O/K2O ratios ranging from 1.48 to 24.33, (with an average of 11.35 ± 12.52). The investigated eclogite samples fall in the picro-basalt and basalt ranges on the total alkalis (K2O + Na2O) vs SiO2 diagram (Fig. S2). The eclogite samples have low LOI values between 0.29 and 1.77 wt%, with an average of 1.20 ± 0.86 wt% (Table S5). The country rock show higher concentrations in SiO2 (70.3 wt%) and K2O (2.52 wt%) than the eclogite samples (Fig. S4), close to the compositions of the upper continental crust (SiO2  66.6 wt%, K2O  2.80 wt%, Rudnick and Gao, 2014), which is consistent with the conclusion that the country rock is derived from continental sediments (Liu et al., 2019a). Generally, the K2O contents of the omphacite included in garnet are much higher (up to 0.7 wt%, Table S7) than the matrix omphacite. 4.2. Potassium isotopes In general, the investigated eclogites exhibit a large spread in d41K values, ranging from
1.64 to
0.24, with an average of
0.76 ± 0.83 (Fig. 2). These values are extremely low compared to those of low-T AOC and also fresh MORB (
0.43 ± 0.17‰, 2sd, Tuller-Ross et al., 2019), with several samples having the lowest d41K values (TB-188 =
1.63 ± 0.04‰, TB-193 =
1.64 ± 0.07‰, TB-197 =
1.39 ± 0.03‰) yet measured in terrestrial

K (mg/g)a Ba (mg/g)b Rb (mg/g)b Cs (mg/g)b Nb (mg/g)b K/Nb Ba/Nb Rb/Nb Cs/Nb d41K

TB-34 TB-169 TB-176 TB-181 TB-182 TB-186 TB-187 TB-188 Duplicateg TB-189 Replicateh TB-190 TB-193 TB-194 TB-195 TB-196 TB-197 TB-198 TB-199 Replicate TB- 200 TB-201 TB-209 TB-183 Replicate TB-182 PHi Duplicate TB-189 PH Duplicate TB-193 PH TB-199 PH Replicate TB-176 OMPi TB-181 OMP TB-199 OMP

1992 1743 1660 1079 1660 2573 1411 830

29.89 41.52 98.30 32.10 86.50 83.20 46.63 21.72

10.22 5.18 3.81 4.09 4.73 7.86 4.07 2.74

0.56 0.35 0.14 0.09 0.11 0.28 0.92 0.14

1.87 0.61 0.90 0.76 1.30 1.01 0.41 7.32

1064 2845 1845 1420 1277 2548 3449 113

5.46 8.45 4.23 5.38 3.64 7.78 9.95 0.37

5.46 8.45 4.23 5.38 3.64 7.78 9.95 0.37

0.30 0.57 0.16 0.12 0.08 0.28 2.24 0.02

2075

54.80

6.89

0.13

1.24

1674

5.56

5.56

0.10

415 581 1328 830 3819 913 9132 747

18.90 7.92 56.45 38.22 80.44 25.64 273.62 22.30

1.08 1.63 3.05 3.20 8.44 2.04 19.82 2.47

0.01 0.03 0.28 0.24 0.53 0.09 0.77 0.03

2.24 0.77 0.73 1.13 3.01 0.59 2.93 1.86

185 755 1818 732 1270 1546 3121 402

0.48 2.12 4.18 2.82 2.81 3.46 6.77 1.33

0.48 2.12 4.18 2.82 2.81 3.46 6.77 1.33

0.00 0.04 0.38 0.21 0.18 0.15 0.26 0.02

1245 581 498 20,920

25.25 13.77 19.58 459.83

5.67 1.47 1.45 87.72

0.18 0.08 0.20 2.48

0.40 1.31 1.24 14.37

3094 445 403 1456

14.08 1.13 1.18 6.10

14.08 1.13 1.18 6.10

0.45 0.06 0.16 0.17


0.56
0.73
0.41
0.76
0.24
0.55
0.34
1.63
1.62
0.70
0.77
0.98
1.64
0.38
0.76
0.28
1.39
0.35
0.89
0.94
0.91
0.86
0.87
0.55
0.52 0.23 0.16
0.41
0.47
0.92
0.98
0.96
1.13
1.10
1.25

95% C.I.c nd

2sd

Dissolutione

Rock type

0.04 0.06 0.05 0.05 0.05 0.06 0.02 0.04 0.06 0.06 0.06 0.07 0.07 0.05 0.05 0.03 0.03 0.02 0.03 0.05 0.03 0.08 0.08 0.05 0.04 0.07 0.05 0.07 0.05 0.06 0.05 0.05 0.05 0.05 0.06

0.10 0.16 0.13 0.14 0.14 0.15 0.05 0.11 0.16 0.15 0.16 0.16 0.16 0.11 0.12 0.07 0.08 0.06 0.07 0.11 0.07 0.16 0.16 0.10 0.07 0.20 0.08 0.19 0.08 0.16 0.14 0.13 0.15 0.15 0.16

Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal Normal High Pressure High Pressure High Pressure

MORB-type eclogite 3 MORB-type eclogite 3 MORB-type eclogite 2 MORB-type eclogite 3 MORB-type eclogite 2 MORB-type eclogite 2 MORB-type eclogite 2 MORB-type eclogite 2 MORB-type eclogite MORB-type eclogite 2 MORB-type eclogite 2 MORB-type eclogite 3 MORB-type eclogite 3 MORB-type eclogite 2 MORB-type eclogite 3 MORB-type eclogite 2 MORB-type eclogite 3 MORB-type eclogite 2 MORB-type eclogite 3 MORB-type eclogite 3 MORB-type eclogite 3 MORB-type eclogite 3 MORB-type eclogite 3 Muscovite quartz schist 2 Muscovite quartz schist 2 Mineral 2 Mineral Mineral 2 Mineral Mineral 2 Mineral 2 Mineral 2 Mineral 3 Mineral 3 Mineral 3 (continued on next page)

9 10 10 10 10 10 9 9 10 9 9 8 8 8 8 8 8 9 8 7 7 6 6 6 5 10 5 10 5 10 10 10 10 10 10

Columnf

H. Liu et al. / Geochimica et Cosmochimica Acta 277 (2020) 206–223

Sample ID

210

Table 1 Potassium and representative trace element concentrations and potassium isotopic compositions of the Sumdo eclogite.

Table 1 (continued) K (mg/g)a Ba (mg/g)b Rb (mg/g)b Cs (mg/g)b Nb (mg/g)b K/Nb Ba/Nb Rb/Nb Cs/Nb d41K

Sample ID i

TB-176Amp TB-181Amp TB-182Amp Duplicate TB-189Amp Duplicate TB-193Amp TB-199Amp


0.50
1.14
0.09
0.11
0.43
0.48
0.66
1.16

95% C.I.c nd

2sd

Dissolutione

Rock type

Columnf

0.05 0.04 0.03 0.04 0.06 0.04 0.05 0.04

0.14 0.11 0.09 0.09 0.17 0.11 0.15 0.11

High High High High High High High High

Mineral Mineral Mineral Mineral Mineral Mineral Mineral Mineral

3 3 3

10 10 10 8 10 10 10 10

Pressure Pressure Pressure Pressure Pressure Pressure Pressure Pressure

3 3 3 H. Liu et al. / Geochimica et Cosmochimica Acta 277 (2020) 206–223

Note: a Calculated from major element, which were determined with an XRF Spectrometer at the laboratory of ALS minerals at Guangzhou. b Trace element concentrations were measured by solution ICP-MS at the University of Science and Technology of China, Hefei (USTC). c 95% confidence interval (c.i.) were calculated from standard deviation of the mean during an analytical session along with correction by Student’s t-factor (95% c.i. = ±t*SD/N1/2). d n: number of analyses. e Normal = The whole rock samples and phengites were dissolved in ultra-pure concentrated HF + HNO3 under heat lamps. High Pressure = The mineral separates were digested into a high pressure teflon bomb and placed in an oven before the ‘‘Normal” dissolution procedural. f Number of column procedures that was carried on the same sample. g Duplicate = repeated measurement of K isotopic ratios on the same purified solutions. h Replicate = repeat column chemistry and instrumental measurement of different aliquots of the same stock solution. i PH = Phengite, OMP = Omphacite, AMP = Amphibole.

211

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Fig. 1. Plots of (a) K (mg/g) vs Rb (mg/g) and (b) K/Nb vs Rb/Nb for the Sumdo eclogite. Data of AOC ODP 417/418 and N-MORB are from Staudigel et al. (1996) and Sun and McDonough (1989), respectively.

samples, excluding the deep-sea pore-fluids (Santiago Ramos et al., 2018, Fig. 2). Interestingly, d41K values of the investigated eclogites are positively correlated with K2O contents and K/Nb ratios (Fig. 2). In contrast, the country rock sample displays similar K isotopic compositions (d41K =
0.54 ± 0.05) to the Bulk Silicate Earth (
0.43 ± 0.17, 2sd, Tuller-Ross et al., 2019; Fig. 2). Similarly, d41K values of the mineral separates are highly variable (
1.25 to +0.23), with phengite ranging from
0.98 to +0.23, omphacite ranging from
1.25 to
1.10, and amphibole ranging from
1.16 to
0.09, respectively (Fig. 3). Overall, the inter-mineral K isotope fractionations between phengite, omphacite and amphibole are significant (Fig. 3). However, the fractionation scales are variable, with D41Kamphibole-phengite and D41Kamphibole-omphacite ranging from
0.04 to +0.63 and from
0.30 to +0.25, respectively (Fig. 3). Notably, phengite show heavier K isotopic compositions than the coexisting omphacite (D41Kphengite-omphacite =+0.25) and amphibole (D41Kphengite-amphibole= +0.03 +0.30, except TB193 is
0.25, Fig. 3). 5. DISCUSSION 5.1. Origin and evolution of the Sumdo eclogite The major, trace element (except K, Rb, Cs, Sr, Ba) and isotope (e.g., Sr, Nd, Pb, O) compositions of the Sumdo

Fig. 2. Plots of (a) d41K (‰) vs K2O (wt%) and (b) d41K (‰) vs K/Nb for the Sumdo eclogite and their country rock. The range of Fresh MORB and Altered MORB are from Tuller-Ross et al. (2019) and Parendo et al. (2017). Note, field for Altered MORB is approximately only, since K concentration (also K/Nb ratio) and d41K data on the same sample are not yet measured. Error bars represent 95% confidence interval. Symbols without error bars have errors that are smaller than the size of the symbol.

Fig. 3. K isotope compositions (‰) of the investigated eclogite rocks and their separated minerals (phengite, omphacite and amphibole). The d41K values of the N-MORB (
0.43 ± 0.17) are after Tuller-Ross et al. (2019). Error bars represent 95% confidence interval. Symbols without error bars have errors that are smaller than the size of the symbol.

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related activities, such as seafloor alteration, prograde metamorphic dehydration and subsequent retrograde rehydration. Therefore, to understand the K isotope characteristics of the Sumdo eclogite, we must consider all processes that could have potentially influenced the evolution of the K isotopic systematics of the Sumdo eclogite samples.

Fig. 4. Plots of K (mg/g) vs 87Sr/86Sr(i) for the Sumdo eclogite. The data of Fresh MORB and AOC ODP 417/418 are from Hofmann (2014) and Staudigel et al. (1996). The sea water Sr isotopic compositions at 300 Ma and today are shown for comparison (Burke et al., 1982).

eclogites demonstrated that they have preserved their premetamorphic signatures, without significantly being affected by eclogite-facies metamorphism (Liu et al., 2019a, 2019b). For instance, the investigated samples have relatively low abundance of incompatible elements and display LREE depleted and HREE flat patterns, suggesting an oceanic crust origin for the protolith of the Sumdo eclogite (Liu et al., 2019a, 2019b; Yang et al., 2009; Zeng et al., 2009). In addition, Sumdo eclogite samples show a positive correlation between initial 87Sr/86Sr ratios (0.703703– 0.706645) and d18O values (+4.9 to +8.9‰), which indicate that the protoliths of the Sumdo eclogite had experienced low-temperature seafloor alteration prior to high-pressure metamorphism (Liu et al., 2019a, 2019b; Zeng et al., 2009). Based on the high Li concentrations (3.64–29.7 lg/g), light Li isotopic compositions (
1.1 to +4.0‰) and Rayleigh dehydration modeling, Liu et al. (2019a) suggested that the Sumdo eclogite have experienced both metamorphic dehydration and retrograde rehydration during early subduction and subsequent exhumation, respectively. This concurs with the conclusion that the eclogite body and its country rocks have undergone amphibolites-facies retrograde metamorphism, simultaneously, which may have rehydrated the Sumdo eclogite during exhumation to the shallower levels (Li et al., 2011a, 2012; Liu et al., 2019b). Furthermore, a comprehensive study of petrography, mineral chemistry and fluid inclusions indicated that the Sumdo eclogite body experienced an initial heating exhumation stage after the peak metamorphic event, during which large amounts of fluids were released by dehydration and preserved as fluid inclusions in metamorphic minerals (Liu et al., 2019b). 5.2. Processes controlling the potassium isotope systematics of Sumdo eclogite Potassium has a high mobility in hydrous fluids, and thus, K isotopes may be strongly fractionated during fluid

5.2.1. Low temperature seafloor alteration During seawater-oceanic crust interactions, temperature clearly affects the direction and magnitude of chemical exchange (Hart and Staudigel, 1982; Kelley et al., 2003; Seyfried and Bischoff, 1979). Potassium is leached from oceanic crust to seawater at high temperatures (>150 °C), but added into the oceanic crust from seawater at low temperatures (<150 °C), which has been explained by temperature dependent ion-exchange reactions between seawater and the alteration products (e.g., smectite, celadonite, Hart and Staudigel, 1982; Kelley et al., 2003; Seyfried and Bischoff, 1979). For instance, enrichment of alkaline elements was observed at depths of about 500 m in profiles of oceanic crust and ophiolite, and the average enrichment factors of K, Rb and Cs in altered basalts from ODP sites 417D, 417A and 418A are 6, 14 and 13, respectively and from ODP site 801 are 4, 9 and 7, respectively (Hart and Staudigel, 1982; Kelley et al., 2003; Staudigel et al., 1996, Fig. 1). Generally, the K enrichment resulted from low temperature alteration is accompanied by increased 87Sr/86Sr ratios and d18O values (McCulloch et al., 1981; Staudigel, 2014; Staudigel et al., 1995). Therefore, the low temperature altered upper oceanic crust represents an alkaline sink (K, Rb and Cs) for the seawater (Hart and Staudigel, 1982; Kelley et al., 2003; Staudigel, 2014; Zhang and SmithDuque, 2014). The investigated Sumdo eclogite samples display generally higher K2O concentrations (0.05–1.10 wt%, with an average of 0.21 ± 0.46 wt%), initial 87Sr/86Sr ratios (0.703703–0.706645) and d18O values (+4.9 to +8.9‰) than the fresh MORB (K2O = 0.1  0.2%, initial 87 Sr/86Sr  0.7028, d18O+5.6‰, (Sun and McDonough, 1989; White and Klein, 2014; Fig. 4). These signatures indicated that the protolith of Sumdo eclogites experienced low temperature alteration, which resulted in the higher K2O concentrations than fresh basalts. However, it is worth noting that the K2O concentrations of the Sumdo eclogites were still lower than the low temperature AOC (ODP 417/418 Super composite samples K2O = 0.56%, Staudigel et al., 1996; ODP 801 Super composite samples K2O = 0.62%, Kelley et al., 2003; IODP 1365 average K2O = 1.51%, Zhang and Smith-Duque, 2014; Figs. 1 and 4). Parendo et al. (2017) investigated ophiolite from the Bay of Islands and observed that the AOC has significantly higher K isotopic compositions (
0.49 to +0.19‰) than the fresh MORB (
0.43 ± 0.17‰, 2sd, Parendo et al., 2017; Wang and Jacobsen, 2016a). Most importantly, a correlation between d41K and 87Sr/86Sr ratios was identified, which demonstrated that the heavier K isotopic compositions of the ophiolite were most likely caused by uptake of seawater K (d41K = +0.02  +0.14) into alteration minerals (Parendo et al., 2017; Wang and Jacobsen, 2016a; Xu et al., 2019). So far, there has been no K isotopic data

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Fig. 5. Plots of (a) d41K (‰) vs Ba (mg/g), (b) d41K (‰) vs Rb (mg/g), (c) d41K (‰) vs Cs (mg/g), (d) d41K (‰) vs Ba/Nb, (e) d41K (‰) vs Rb/ Nb, and (f) d41K (‰) vs Cs/Nb for the Sumdo eclogite and their country rock. The range of Fresh MORB and Altered MORB are from Tuller-Ross et al. (2019) and Parendo et al. (2017). Error bars represent 95% confidence interval. Symbols without error bars have errors that are smaller than the size of the symbol.

reported for in situ low temperature altered oceanic crust samples, except an abstract, which suggested that due to uptake of isotopically heavy seawater K (d41K = +0.02  +0.14), AOC could have heavier K isotopic compositions relative to fresh MORB. In summary, the protoliths of the Sumdo eclogite samples should have high K2O concentrations and high K isotopic compositions.

5.2.2. Prograde metamorphic dehydration During the subduction of oceanic crust, incompatible elements (e.g., large ion lithophile elements, LILE) and water are partially recycled in fluids to the hanging wedge mantle, ultimately producing the formation of arc magmas (Bebout, 2014; Becker et al., 2000; Schmidt and Poli, 2014). Becker et al. (2000) investigated a range of incompatible to

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Fig. 6. Plots of (a) Li vs K, (b) d41K vs Li, (c) d7Li vs K and (d) d41K vs d7Li for the Sumdo eclogite and their country rock. The Li and K concentration of the N-MORB are from Sun and McDonough (1989). The d41K range of Fresh MORB and Altered MORB are from TullerRoss et al. (2019) and Parendo et al. (2017). The Li isotopic data of Sumdo eclogite are from Liu et al. (2019a).

compatible trace elements in high-pressure metamorphic rocks and observed that the most K, Rb and Ba appears to be lost at temperatures < 600–700 °C, while little Th, Ti and Zr (<10-20%) were lost and nearly all Nb and Zr were retained in the eclogites. Although Nb and Th have similar incompatibility to K and Rb, they show much more conservative behavior during fluid-rock interactions. For this reason, these element ratios (e.g., K/Nb, K/Th, Rb/Nb, Rb/Th) can be used to assess the losses or gains of LILE (e.g., K, Rb, Cs, Bebout, 2014; Becker et al., 2000; Chen et al., 2018; Cao et al., 2019). The Sumdo eclogite samples have K/Nb (113–3449, with an average of 1500 ± 2090) and K/Th (900–97487, with an average of 15553 ± 42181) ratios that are comparable to those of fresh MORB, but much lower than those of the altered MORB (e.g., AOC, ODP 417/418, K/Nb = 700–24000, with an average of 7700 ± 14000; K/Th = 13000–430000, with an average of 130000 ± 250000, Staudigel et al., 1996, Figs. 1 and S2). And, the clear positive correlation is the outstanding signature of the K/Nb vs K/Th diagram (Fig. S3a). The Rb/Nb vs Rb/Th diagram shows a similar trend to that in the K/Nb vs K/Th diagram, indicating the similar geochemical behavior of K and Rb during subduction dehydration (Becker et al., 2000; Zack et al., 2001, Fig. S3).

Compared to their protoliths, namely the low temperature AOC, the Sumdo eclogites have K/Nb, K/Th, Rb/Nb and Rb/Th ratios that are lower by 1 to 2 orders of magnitude, indicating a large loss of K and Rb from their protolith (Fig. S3). In addition, a K-87Sr/86Sr diagram can also be used to estimate the K loss (Becker et al., 2000). The difference between the correlation of altered MORB and the Sumdo eclogites (different slopes) demonstrate a strong depletion of K (Fig. 4). Notably, the Sr isotopic composition of seawater at 300 Ma is much lower than that of the modern seawater (Burke et al., 1982) and the87Sr/86Sr(i) of Sumdo eclogites nearly match that of the seawater at 300 Ma (Fig. 4), which indicates that the protoliths of the Sumdo eclogites should be even more enriched in K when compared to modern altered MORB (e.g., ODP 417/418 in Fig. 4). Thus the Sumdo eclogites have experienced even stronger depletion of K than they appear in Fig. 4. The strong depletion of K, Rb and Cs relative to similar incompatible elements such as Nb and Th in Sumdo eclogite is also obvious in N-MORB normalized spider diagrams (Fig. S4). All these observations above suggest that K and Rb are strongly depleted during the metamorphic dehydration, which is consistent with other high-pressure metamorphic rocks from a number of high-pressure ter-

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ranes (Bebout et al., 2007; Becker et al., 2000; Zack et al., 2001). The d41K values of the investigated samples are exceptionally light compared to their protoliths (i.e., low-T AOC), fresh basalts, and almost all other terrestrial rock types that have been studied so far, and are therefore unlikely to simply reflect the K isotopic compositions of their protoliths (Li et al., 2019a, 2016b; Morgan et al., 2018; Parendo et al., 2017; Santiago Ramos et al., 2018; Wang and Jacobsen, 2016a). In addition, the positive correlations between d41K values and K2O contents (also Ba, Rb, Cs, K/Nb, Ba/Nb, Rb/Nb, Cs/Nb) suggest that metamorphic dehydration may explain the obviously low d41K values in the investigated eclogite samples (Figs. 2 and 5). As evidenced by Santiago Ramos et al. (2018), the fractionation during partitioning of K isotopes between mineral and fluids in rapidly accumulating deep-sea sedimentary systems is influenced by changes in cation coordination number with the heavy isotope of an element tending to be enriched in phases with the strongest bonds and lowest coordination (Schauble, 2004). Aqueous K has a coordination number of 6 ± 2, therefore, we suggest during dehydration, heavy K isotopes preferentially enter into the fluid (low coordination number) while the light K isotopes remain in silicate minerals (high coordination number, Santiago Ramos et al., 2018; Varma and Rempe, 2006). Below (Discussion Section 5.3), we will use a Rayleigh fractionation model to simulate the K isotopic systematic compositions of the Sumdo eclogite, which again demonstrates that the low d41K values were most likely caused by dehydration during subduction. 5.2.3. Retrograde rehydration The recent Li isotopic investigations on the same samples observed that the Sumdo eclogites have comparable Li concentrations (3.64–29.7 lg/g), but lower d7Li values (
1.1 to +4.0‰) when compared with their protolith (Liu et al., 2019a). Based on a two-stage modeling, Liu et al. (2019a) suggested that a single stage of dehydration cannot explain their Li isotopic characteristics, but retrograde rehydration (fluid < 5%) is necessary to obtain the high Li concentration and low Li isotopic compositions of the Sumdo eclogite rocks (Fig. 6, Liu et al., 2019a; Marschall et al., 2007; Penniston-Dorland et al., 2010; Simons et al., 2010). In addition, the fluid exchange between the retrograde fluids (<5%) and the eclogite body can be documented by the relative enrichment of Ba and Cs compared with K and Rb in the Ba-K and Cs-K diagrams and the N-MORB normalized diagram (Fig. S4). Therefore, such fluid exchange could also influence the K isotopic systematics of the Sumdo eclogites. However, no correlation between K/d41K values and Li/d7Li values have been observed, which indicates that the retrograde (fluid < 5%) rehydration is too weak to affect the K isotopic compositions of the investigated samples and has no resolvable effects (Fig. 6). The modeling shows that Li was lost during dehydration, while during exhumation the infiltration of sediment-derived aqueous fluid can add Li to the dehydrated residual slab (See the dehydration and rehydration trends in Fig. 6, Liu et al., 2019a). In Fig. 6a-b, some

Sumdo eclogites show higher Li concentrations than the AOC (rehydration caused), however, seldom samples have higher K concentrations and d41K values than the AOC, indicating that rehydration has a limited effect on K isotopic systematics. In addition, dehydration caused a continuous decrease in d7Li values from +7.5‰ to +4.5‰, which are still higher than the d7Li values of Sumdo eclogites (Liu et al., 2019a). Therefore, a rehydration stage during exhumation is necessary to reduce the d7Li values from +4.5‰ to +1.9‰ (Liu et al., 2019a). However, as shown in Fig. 6c-d, during the rehydration stage, the d7Li values of Sumdo eclogites display no correlation with K and d41K values, which generally remain constant. The decoupling between K (Rb) and Li (Ba, Cs) can be explained by the fact that Li (Ba, Cs) displays stronger fluid mobility than K (Rb), therefore, retrograde fluids could easily alter Li (Ba, Cs) abundances, without having as significant of an effect on K (Rb, Becker et al., 2000; Cao et al., 2017c; Du et al., 2018, 2019; Liu et al., 2019a; Zack et al., 2001). Furthermore, the recent first-principles calculations inferred that the equilibrium K isotope fractionations between K-bearing solutions and most silicate minerals would enrich 41K in the fluid (Li et al., 2019b; Zeng et al., 2019). Therefore, equilibrium fractionation between fluid and schists should result in the enrichment of K2O and elevation of the K isotopic composition (Morgan et al., 2018; Li et al., 2019a). However, the investigated samples have generally lower K2O contents (0.21 ± 0.46 wt% vs 2.52 wt%) and are isotopically lighter than the schists (-0.76 ± 0.83‰ vs
0.55‰), which suggest that retrograde rehydration is unlikely to affect the K isotopic compositions of the Sumdo eclogite (Figs. 2 and 5). 5.2.4. Diffusion-driven fractionation of K isotopes Laboratory experiments have demonstrated that kinetic stable isotope fractionations can also occur during diffusion even at high temperatures (Huang et al., 2010; Richter et al., 2014, 2003). For instance, thermodiffusion (Soret effect) experiments have showed that in a temperature gradient, K isotopes can fractionate by approximately 1.06‰ per 100 °C temperature difference, with the hot end of the gradient becoming relatively enriched in lighter K and the cold end relatively enriched in heavy K (Richter et al., 2014, 2003). A study on the reverse silicate weathering shows that the diffusive K fractionation plays an important role in determining the d41K value of deep-sea pore-fluids and authigenic marine silicates (Santiago Ramos et al., 2018). In particular, the K isotopic compositions of porefluids from slow-accumulating sedimentary systems are mainly controlled by K fractionation during diffusion (Santiago Ramos et al., 2018). However, due to the faster diffusion rate of the lighter isotope compared to the heavier isotope, diffusion processes should lead to higher K concentrations covarying with lighter isotopic compositions. For instance, if the K of the country rock was diffused into the Sumdo eclogite body, we should observe higher K concentrations in the sample correlating with lighter isotopic compositions. However, d41K values and K2O contents (also K/Nb ratios) of Sumdo eclogites show clear positive correlations, namely the lower

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the K concentrations with the lighter the K isotopic compositions (Figs. 2 and 5), which indicate that diffusion can not explain the K isotopic systematics of the Sumdo eclogites. 5.3. Potassium isotopes fractionation during dehydration Here we model the K isotope fractionation during subduction dehydration by Rayleigh Fractionation. Three parameters need to be constrained: (1) the K concentration and K isotopic composition of the protolith, (2) the fractionation factor between the residual slab and the released fluid, and (3) the partition coefficient between the residual slab and the released fluid. Previous investigations have demonstrated that the Sumdo eclogite represents the product of subducted lowT AOC; therefore, the K concentrations and isotopic compositions of the low-T altered MORB or ophiolite are used to represent the compositions of its protolith (Liu et al., 2019a, 2019b; Parendo et al., 2017; Zeng et al., 2009). The previous studies of K isotopes in dissolved loads and sediments from rivers suggested that the heavy K isotope (41K) is preferentially partitioned into aqueous solutions, whereas the light K isotope (39K) is preferentially incorporated into secondary minerals, resulting in the higher d41K values of riverine dissolved loads (Li et al., 2019a). These observations in rivers are consistent with the ab initio calculations (Li et al., 2019b; Zeng et al., 2019), where aqueous solutions are enriched in heavier K isotopes than most silicate minerals, such as illites. In contrast, a recent study on hydrothermal alteration in a porphyry system revealed that K isotope fractionation behaviors at elevated P-T conditions can be different from those at surface weathering conditions (Li et al., 2020). And, they attributed the increase in d41K in altered rocks to hydrothermal alteration, which caused preferential incorporation of heavy K isotope in alteration products, particularly muscovites (Li et al., 2020). In detail, they assumed that KCl ion pair in hightemperature, high-salinity hydrothermal fluid enriches in light K isotope, thus the high d41K characteristics of the altered rocks can be explained by equilibrating with such fluids (Li et al., 2020). However, the direction and degree of K isotopic fractionations between K-bearing solutions and minerals are depending on the K speciation in the hydrothermal fluids and coexisting minerals (Li et al., 2019b; Zeng et al., 2019). Thus considering the mass balance and the fact that most of the Sumdo eclogite samples display extremely light K isotopes, we hypothesized that 41 K will tend to enter the fluid during metamorphic dehydration and the K isotope fractionation factor (a = (41K/39K)water/(41K/39K)mineral) should be higher than 1 (Li et al., 2019a). In this study, due to the lack of experimental data, we simply assume that the a value between fluid and residual slab remains constant during subduction and the best fitted value (1.0015) is adopted (Table S8). Recently, a > 1‰ range in d41K values -from
0.11‰ to
1.36‰- is observed in a suite of igneous and metamorphic pegmatites (including K-feldspars:
0.32‰ to
1.19‰, micas:
0.11‰ to
1.36‰, and astrophyllite:
0.71‰), which clearly shows that a fluid system can facilitate K isotopic fractionations even in high temperature igneous and

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metamorphic conditions (Morgan et al., 2018). Therefore, such a a value (1.0015) is reasonable. Phengite (K2O up to 11 wt%) and sometimes amphibole (K2O up to 0.2 wt%) are the dominant hosts for K in eclogite and incorporate more than 90% of the budgets of K in the whole rocks (Liu et al., 2019b). Therefore, the volume content of phengite is expected to significantly influence the bulk partition coefficient for K (DMin/Fluid) between the residual slabs and dehydrated fluids, with lower content resulting in lower D value. The variable bulk D values (0.09–0.26) are calculated based on the estimated volume contents of phengite, amphibole, and omphacite and their mineral partition coefficients (4.7, 0.14 and 0.002, respectively, Table S8, Liu et al., 2019b; Zack et al., 2001). The water contents of the protolith were assumed to be 6% and 8% for slightly and highly altered AOC, respectively, according to Marschall et al. (2007) and Liu et al. (2019a). Based on these parameters, Rayleigh Fractionation was modeled to calculate the K concentration and d41K values of the residual slabs and instantaneous fluid in equilibrium with the residual slabs, using the following equations: C solid ¼ C 0  ð1
F Þ1=D
1

ð1Þ

d41 K solid ¼ ½ðd41 KÞi þ 1000  f ða
1Þ
1000

ð2Þ

C fluid ¼ C 0 =D  ð1
F Þ1=D
1

ð3Þ

d K fluid ¼ a  ½ðd KÞsolid þ 1000
1000

ð4Þ

41

41

41

where C0 and (d K)i are the initial concentration and isotopic composition, F is the fraction of liquid relative to the solid, and f is the ratio of Csolid/C0. D is the bulk partition coefficient, and a is the fractionation factor. Different degrees of water loss were modeled in this study, and the results show that the dehydration is characterized by a continuous decrease of K concentration (535–9774 mg/g) and d41K value (
1.37 
0.13 ‰) of the residual slabs (Table S8, Fig. 7). Accordingly, the model indicates that the subduction can account for a 47 % loss of the K budget in the rocks and a decrease of 0.94‰ in d41K values (Table S8, Fig. 7). It should be noted that the modeling results suggest that the less a basalt has been altered (through low-T alteration), corresponding to lighter (d41K)i, the lighter this rock becomes after metamorphic dehydration. This is similar to the case of Li and B isotopes (Bouvier et al., 2010; Liu et al., 2019a; Marschall et al., 2007; Simons et al., 2010). In other words, the K isotopic compositions of high-K eclogites (protoliths were more altered) were not significantly changed by dehydration reactions during subduction, which is consistent with the observation that the original LILE patterns of high-K eclogites (e.g., Trescolmen, Zermatt and Confin) are still preserved after metamorphic dehydration (Becker et al., 2000; Zack et al., 2001). As shown in Fig. 7, the dataset of Sumdo eclogite can be reproduced by these calculations in terms of (1) their lower K concentrations and lighter d41K values compared to the low-T AOC, and (2) a clear positive correlation between the K concentrations and the d41K values (Fig. 7).

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Fig. 7. Modeling results for K isotopes dehydration of the low-T altered oceanic crust. The cyan and yellow areas represent the fields for the assumed low-T altered MORB (Parendo et al., 2017) and the residual slab after completely dehydration, respectively. The blue five-pointed stars indicate the K concentration and d41K values used for the Rayleigh fractionation modeling. Pink shaded fields include the Sumdo eclogite samples. Rayleigh Fractionation was modeled according to Eqs. (1) and (2) with constant a = 1.0015 and variable D = 0.09–0.26 (see text for details). Solid lines with tick marks (numbers next to the marks are the proportions of water loss) denote Rayleigh dehydration trends of the altered MORB. The residuals with the same degree of dehydration are linked by dash lines.

In contrast, the released fluids (K = 5944–37592 mg/g, d41K = +0.13  +1.37‰) display higher K concentrations and are isotopically heavier than the residual slab and Bulk Silicate Earth (Table S8). This is substantiated by the systematic increase in K concentrations (86–587 mg/g) of the pore water collected from the Mariana mud volcanoes as a function of the distance from the trench (55–86 km) and the high K concentrations in arc lavas (e.g., Kamchatka arc, Mariana arc, Churikova et al., 2001; Fryer et al., 2018; Hulme et al., 2010; Ishikawa and Tera, 1999). Thus, it is very likely the arc volcano would be enriched in heavier K isotopes. However, no samples of arc lava have been analyzed for K isotopes and this hypothesis needs to be further tested. As such, we conclude that the extremely low d41K values observed in the Sumdo eclogite may reflect the metamorphic dehydration during the subduction of low-T altered oceanic crust. 5.4. Inter-mineral potassium isotopic fractionation The separated minerals from the Sumdo eclogite display a wide range in K isotopic compositions (
1.25 to +0.23‰, Fig. 3). Generally, phengites show heavier K isotopic compositions than the coexisting omphacite (D41Kphengite-omphacite = +0.25‰) and amphiboles (D41Kphengite-amphibole = +0.03+0.30‰, except TB193 is
0.25‰), which, in the first order, may suggest equilibrium K fractionation between coexisting minerals, as phengite has lower K coordination than omphacite and amphibole (phengite, 6; omphacite, 7 to 8 and amphibole, 8; Li

et al., 2019b; Harlow George, 1997; Hawthorne and Calvo, 1977; Huebner and Papike, 1970; Fig. 3). Ab initio calculations show that the K-bearing minerals in which K has a lower coordination number (stronger bond) would be enriched in heavy K isotopes (Li et al., 2019b; Zeng et al., 2019). Therefore, the inter-mineral fractionation of K isotopes is consistent with the ab initio calculation. However, the observed D41Kamphibole-phengite and D41Kamphibole-omphacite values for the six samples are variable and do not display correlations with any other geochemical parameters (Table 1), which suggests disequilibrium K isotope fractionations between phengite and omphacite/amphibole. Mineral chemistry investigations show that the phengites display decreasing Si contents from the core (Si pfu: 3.52–3.59) towards the rim (Si pfu: 3.38–3.44), indicating that the phengites in the Sumdo eclogite were affected by retrograde metamorphism (Liu et al., 2019b). This is consistent with the replacement of omphacite by amphibole ± plagioclase (Liu et al., 2019b). These observations, together with the multiple generations of amphiboles identified by petrography observations, suggest that disequilibrium K isotope fractionations may result from retrograde metamorphism or K exchange between different mineral growths. However, it should be noted that the whole rock K isotopic compositions are not in mass balance with the mineral separates analyzed for K isotopic compositions (Fig. 3). For samples TB-176, TB-181, and TB-199, the d41K values of bulk rocks are all higher than that of the analyzed mineral separates, while for samples TB-182, TB-189, and TB193, the d41K values of the bulk rocks are all lower than that of the analyzed minerals. To better explain the disequilibrium K isotope fractionations, we calculated the K mass balance by the mineral volume contents and K concentrations to evaluate whether the K in the bulk rocks can be balanced solely by the minerals (Table S1). The results indicated that the phengites are the dominant K host of the bulk rock in all the samples (Table S1). The higher the phengite volume, the higher the K concentrations of the whole rock. Therefore, we infer that the mismatch between mineral separates and whole rock K isotopic values for samples TB-176 and TB-181 result from the lack of K isotopic measurements of phengites (Fig. 3). This is also consistent with the observed heavy K isotopic compositions of phengite in samples TB-182, TB-189 and TB-199 (Fig. 3). Similarly, it is reasonable for sample TB-199 that the whole rock (d41K =
0.91 ± 0.05‰) and phengite (d41K =
0.97 ± 0.05‰) have indistinguishable K isotopic compositions, even though both omphacite and amphibole have slightly lower d41K values (Fig. 3, Table 1). For samples TB-182, TB-189 and TB-193, the whole rock K isotopic compositions are lighter than the phengite, which means that there must be missed phases that can host light K isotopes (Fig. 3). Previous investigations have observed large amounts of mineral (e.g., omphacite) and fluid inclusions hosted by porphyroblast garnets and omphacites, respectively (Liu et al., 2019b). Specifically, the omphacite inclusions hosted by garnets have much higher K concentrations (up to 0.7 wt%, Table S7) than

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Fig. 8. Schematic model of K isotope systematics during subduction and relative K inventory. Blue (1) and red (2) arrows indicate that K is taken up by basalts during low-temperature alterations, while leached from the basalts during high-temperature hydrothermal alteration. During seawater alteration, due to the uptake of isotopically heavy K (d41K  +0.02 to +0.14‰) from seawater, the AOC display heavier K isotopic compositions than the Fresh MORB (d41K 
0.43 ± 0.13‰, Parendo et al., 2017; Wang and Jacobsen, 2016). The subducted crust will release K, characterized by isotopically heavy component, into the above mantle wedge, which may produce anomalous high d41K values in the arc lavas (3). Meanwhile, if the residual slab reaches the deep mantle, then the extremely light K will be delivered into the mantle (4) and might eventually be sampled by plume magmatism (5). The range of Fresh MORB, Altered MORB and continental crust are from TullerRoss et al. (2019), Parendo et al. (2017) and Huang et al. (2019). The framework of this diagram is after Hanyu et al. (2019). (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

the matrix omphacite, which is comparable with the K-rich diopside (up to 1.5 wt%) included in garnets from UHP metamorphic rocks from Kazakhstan (Sobolev and Shatsky, 1990) and experimentally produced K-rich pyroxene (up to 1.5 wt%, Harlow George, 1997). This together with the low d41K values of omphacites from samples TB-176, TB-181, and TB-199 strongly suggest that the omphacite inclusions in garnet are the most likely candidate to host light K isotopes (Fig. 3). To better evaluate this explanation, K isotopic mass balance was calculated by assuming appropriate K isotopic compositions of omphacite (Table S1), the results show that for samples TB-182 and TB-189, if the d41K values of the omphacite were 0.5–1.0 per mil lighter than the measured whole rocks, then less than 5% of included omphacites are adequate to balance the whole rock K isotopic compositions. These amounts are consistent with the petrographic observations. However, for sample TB-193, even estimated d41K value of the included omphacite is 1.5 per mil lighter than the measured whole rock, more than 10% of included omphacite is needed to balance the whole rock K isotopic composition. We have to admit that this amount is much larger than the estimated volume of the included omphacite based on the petrographic observation. At such condition, the fluid inclusion is necessary to explain the whole rock K isotopic mass balance issue. As a fluid mobile element, K is enriched and mobile in fluids. Therefore, the fluid inclu-

sions (NaCl dominated, 10–22 wt% NaCl, Liu et al., 2019b) released during dehydration of the subducting oceanic crust may be the other process to account for the mismatch between mineral separates and whole rocks for K isotope compositions. However, limited by the analytical method, we acknowledge that the K isotopic compositions of the mineral and fluid inclusions cannot yet be obtained and more data on mineral separates are desirable before we can fully reveal the inter-mineral potassium isotopes behavior. 5.5. Implications for potassium recycling Schmidt (1996) performed a series of experiments on MORB to constrain the recycling of K in subduction zones. These experiments demonstrated that phengite, and at higher pressure K-hollandite (KAlSi3O8) and also omphacite (an experiment produced fictive high K component, KAlSi2O6, Harlow, 1997), can transfer K to great depths (up to 300 km) in subduction zones (Schmidt and Poli, 2014; Schmidt, 1996). At shallow levels (<50 km), the descending oceanic crust only loses a small amount of K due to the breakdown of amphibole, which was confirmed by the limited loss of K in the forearc zone of Mariana subduction (Fryer et al., 2018; Hulme et al., 2010; Poli and Schmidt, 1995; Schmidt and Poli, 2014; Schmidt, 1996). For depths of 100 to 300 km, K-omphacite might form continuously from phengite through the reaction

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phengite = K-omphacite + enstatite + coesite + K-rich fluid, resulting in the continuously release of K from the subducted crust (Schmidt and Poli, 2014; Schmidt, 1996). When the wet solidus of K-bearing MORB was achieved with increasing P and T (1050 °C at 9.0 GPa), the descending oceanic crust together with metasediments, may generate a pulse of K-rich fluid, which was evidenced by the subsolidus reaction phengite + coesite + omphacite = garnet + melt (Schmidt and Poli, 2014; Schmidt, 1996). This conclusion is substantiated by the behavior of K in eclogite from Trescolmen, Central Alps (Zack et al., 2001). However, investigations of high P metamorphic oceanic basalts suggested that a large fraction (>90%) of K (also Rb and Ba) was lost at temperatures < 600–700 ° C (Becker et al., 2000; Tatsumi, 1989). Our model results suggest that less than 47% of K was released to the overlying mantle wedge during subduction (<100 km), which is similar to Schmidt (1996)’s experiments and Zack et al. (2001)’s observation (39% to 48% loss), but much lower than Becker et al. (2000)’s result. Significantly, after dehydration high-K eclogites (namely the strongly altered AOC) still preserve higher K concentrations than the depleted mantle, which, therefore, may subsequently cause K heterogeneity within the mantle (Hofmann, 2014; Schmidt and Poli, 2014; Schmidt, 1996; Sun et al., 2008). Interestingly, our research reveals that subduction dehydration can cause large K isotopic fractionation, during which an isotopically heavy K component was released into the mantle wedge, while the light component can be delivered to the deep mantle by subduction (Table S8, Fig. 8). Therefore, it is reasonable to speculate that the island arc lavas, derived from the mantle wedge metasomatized by slab-derived fluid, will display anomalously high d41K values (Fig. 8). On the other hand, if the Sumdo eclogite is broadly representative of subducted oceanic crust, then when the slab reaches the deep mantle its K isotopic composition will be significantly lighter than that of the depleted mantle (Wang and Jacobsen, 2016a; Tuller-Ross et al., 2019). However, one needs to be careful when making such assertions, because that while low-K eclogite (protoliths were less altered through low-T hydrothermal alteration) can be enriched in extremely low K isotope compositions, the high-K eclogite (protoliths were more altered through low-T hydrothermal alteration) would have a K isotope composition similar to the mantle. Thus, the extremely low K isotope signature of eclogite would be diluted and less evident in the mantle. This suggestion can be supported by a recent global systematic study on oceanic basalts Tuller-Ross et al. (2019), where no heterogeneity of K isotopes found among mantle-derived samples. In addition, based on the enrichment of Nb, Ta and high Nb/ U ratios, Sun et al. (2019) indicated that the heavy K isotopic compositions of the sodic lavas resulted from the incorporation of highly altered oceanic crust in their source region, indicating that the highly altered oceanic crust may still preserve the heavy K isotopic compositions. Nevertheless, this study shows that some subducted low-T AOC may contribute light K isotopic crustal material to the mantle, which might eventually be sampled by

plume magmatism and is likely to affect the long-term evolution of the mantle (Fig. 8, Hofmann, 2014; Sun et al., 2008). 6. CONCLUSIONS

(1) The Sumdo eclogites have d41K values ranging from
1.64 to
0.24‰, significantly lower than that of their protolith, namely the Low-T AOC, and the fresh MORB. This, together with the positive correlation between d41K values and K concentrations (also Ba, Rb, Cs, K/Nb, Ba/Nb, Rb/Nb, Cs/Nb), suggest that K isotopic fractionation was most likely caused by subduction dehydration during prograde metamorphism. (2) The K isotope characteristics of the Sumdo eclogites can be reproduced by Rayleigh fractionation with a best fitted fractionation factor (a = 1.0015) and variable bulk partition coefficient (D = 0.09–0.26). The modeling results show that an isotopically heavier K component was released into the mantle wedge, while the lighter component can be delivered to the deep mantle by subduction. (3) This investigation suggests that dehydrated fluid bearing isotopically heavy K will metasomatize the above mantle wedge, which may generate island arc lavas characterized by anomalous high d41K values, while the deeply subducted oceanic crust might introduce crustal materials with light K isotopic compositions into the mantle and might eventually be sampled by plume magmatism.

Declaration of Competing Interest The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper. ACKNOWLEDGEMENTS We acknowledge Dr. Heng Chen and Zhen Tian for their assistance during the K isotopes measurement. Yangyang Wang, Haihao Guo and Lingsen Zeng are thanked for their help in the field work. We thank Dr Fang-Zhen Teng and three anonymous reviewers for thorough and helpful reviews that greatly improved the manuscript. The editor Jeffrey Catalano is acknowledged for efficient editorial handling. This study is financially supported by the Strategic Priority Research Program (B) of the Chinese Academy of Sciences (XDB18020102), Natural Science Foundation of China (41903006), the Taishan Scholar Program of Shandong (ts201712075), AoShan Talents Cultivation Program Supported by Qingdao National Laboratory for Marine Science and Technology (2017ASTCP-OS07) and Qingdao National Laboratory for Marine Science and Technology (2017ASKJ02). Haiyang Liu is partly supported by the China Postdoctoral Science Foundation (2019M652497), the postdoctoral innovation project of Shandong province and the postdoctoral applied research program of Qingdao City to Haiyang Liu. Kun Wang and Brenna Tuller-Ross were supported by the McDonnell Center for the Space Sciences.

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