# Constraining the behavior of gallium isotopes during evaporation at extreme temperatures

Josh Wimpenny, Naomi Marks, Kim Knight, Lars Borg, James Badro,
Frederick J. Ryerson

► To cite this version:

Josh Wimpenny, Naomi Marks, Kim Knight, Lars Borg, James Badro, et al. Constraining the behavior of gallium isotopes during evaporation at extreme temperatures. Geochimica et Cosmochimica Acta, 2020, 286, pp.54-71. 10.1016/j.gca.2020.07.006. hal-03121889

HAL Id: hal-03121889

https://hal.science/hal-03121889v1

Submitted on 5 Aug 2025

HAL is a multi-disciplinary open access
archive for the deposit and dissemination of sci-
entific research documents, whether they are pub-
lished or not. The documents may come from
teaching and research institutions in France or
abroad, or from public or private research centers.

L'archive ouverte pluridisciplinaire HAL, est
destinée au dépôt et à la diffusion de documents
scientifiques de niveau recherche, publiés ou non,
émanant des établissements d'enseignement et de
recherche français ou étrangers, des laboratoires
publics ou privés.

<figure><img src="image_1.png" /></figure>

Distributed under a Creative Commons Attribution 4.0 International License
<figure><img src="image_2.png" /></figure>

# Constraining the Behavior of Gallium Isotopes During Evaporation at Extreme Temperatures

J. Wimpenny, N. Marks, K. Knight, C. Sio, L. Borg, J. Badro, F. Ryerson

December 5, 2019

Geochimica Et Cosmochimica Acta
This document was prepared as an account of work sponsored by an agency of the United States government. Neither the United States government nor Lawrence Livermore National Security, LLC, nor any of their employees makes any warranty, expressed or implied, or assumes any legal liability or responsibility for the accuracy, completeness, or usefulness of any information, apparatus, product, or process disclosed, or represents that its use would not infringe privately owned rights. Reference herein to any specific commercial product, process, or service by trade name, trademark, manufacturer, or otherwise does not necessarily constitute or imply its endorsement, recommendation, or favoring by the United States government or Lawrence Livermore National Security, LLC. The views and opinions of authors expressed herein do not necessarily state or reflect those of the United States government or Lawrence Livermore National Security, LLC, and shall not be used for advertising or product endorsement purposes.
1 Constraining the Behavior of Gallium Isotopes During Evaporation at
2 Extreme Temperatures

3 Josh Wimpenny¹, Naomi Marks¹, Kim Knight¹, Lars Borg¹, James Badro², Rick Ryerson¹

4 ¹Lawrence Livermore National Laboratory, Livermore, CA 94550, USA

5 ²Université de Paris, Institut de physique du globe de Paris, CNRS, 75005 Paris, France

7 29-June-2020

8 Geochimica Et Cosmochimica Acta

9 LLNL-JRNL-798685

10

11
12 Abstract

13 Renewed interest in gallium isotope systematics has stemmed from the fact that Ga is
14 moderately volatile and is hypothesized to undergo kinetic fractionation during evaporation.
15 Here, we present the first Ga isotope data from terrestrial volatile depleted samples including a
16 suite of experimentally heated rhyolitic soils, fallout melt glass, and splash-form tektites from
17 the Australasian strewn field (hereafter termed australite tektites). The Ga in these samples is
18 isotopically heavy compared to Ga in terrestrial basalts and estimates for the composition of
19 the bulk silicate Earth (BSE). For each sample suite the isotopic fractionation of Ga scales with
20 the degree of Ga depletion, consistent with isotopic fractionation caused by evaporation.

21 The rapid experimental heating of rhyolitic soil to temperatures ranging between 1600-2200 °C
22 resulted in volatile loss from the starting soil. Based on the fraction of Ga that was evaporated
23 and the degree of Ga isotopic fractionation between starting soil and experimental samples, we
24 calculate a fractionation factor (α) of 0.99891 ± 0.00024. This is within uncertainty of the
25 fractionation factor we previously calculated for Zn isotopes in the same sample suite (0.99879
26 ± 0.00013). Although Ga isotopic data from nuclear fallout melt glass is less coherent, the Ga
27 isotope systematics are generally consistent with a suppressed fractionation factor of
28 approximately 0.9995-0.9998 during evaporation, which is also similar to the behavior of Zn
29 systematics. Thus, although the fractionation factors obtained from the laser heating
30 experiments and fallout melt glass are different, in both cases Ga and Zn behave similarly, as
31 evidenced by the covariation of δ<sup>71</sup>Ga and δ<sup>66</sup>Zn in these samples.

32 The behavior of Ga isotopes in australite tektites is more difficult to constrain because we do
33 not know the location of the impact site and hence the chemical composition of the target
34 rocks. Nevertheless, based on the composition of more volatile rich Muong-Nong type tektites,
35 we estimate that evaporative fractionation of Ga occurs with an α between 0.9998 and 0.9987;
36 broadly consistent with data from the laser heating experiments and nuclear fallout glass.
37 There is no correlation between δ<sup>71</sup>Ga and δ<sup>66</sup>Zn values in australite tektites which is likely to
38 reflect inherited isotopic heterogeneity from weathered precursor material in combination with
39 varying extents of evaporative loss during tektite formation.

40 Gallium isotope ratios in mare basalts are generally isotopically heavy compared to basalts from
41 Earth. Individual mare basalts have δ<sup>71</sup>Ga and δ<sup>66</sup>Zn values that do not correlate, contrary to
42 data from the laser levitation experiments and nuclear fallout glass. This suggests that δ<sup>71</sup>Ga
43 and/or δ<sup>66</sup>Zn values were fractionated by geologic processes after the Moon had accreted.

44

45 1. Introduction
Gallium (Ga) is a trace element in the Earth's crust with typical concentrations of ~17 ppm (Rudnick & Gao, 2003). Although it is considered a siderophile element, it has variable siderophile characteristics in natural systems; being relatively abundant in iron meteorites (> 100 ppm) while it is modeled to be scarce in Earth's core (Blanchard et al., 2015; Humayun 2017). In the crust and mantle Ga behaves similarly to aluminium, as it is relatively incompatible in the ferromagnesian minerals olivine and pyroxene, while it is relatively compatible in feldspars and spinel (Malvin & Drake 1987; Le Roux et al., 2015). Gallium is a moderately volatile element (MVE) with a 50% condensation temperature of 695 °C (Lodders 2003). Recent studies have shown that the stable isotopes of MVEs including Ga, Zn, Cu, Rb, and K are highly sensitive tracers of volatile loss associated with geologic processes (e.g. Moynier et al., 2009; Herzog et al., 2009; Wang & Jacobsen 2016; Kato & Moynier 2017; Pringle & Moynier 2017; Wimpenny et al., 2019). During evaporation the light isotopes are preferentially concentrated in the vapor phase whereas the residual phase is correspondingly enriched in the heavy isotopes. This property, coupled with the observation that stable isotope fractionation is negligible in magmatic systems (e.g. Chen et al., 2013; Kato et al., 2017), has led to the use of MVEs as tracers of evaporative processes on volatile depleted bodies, such as the Moon (Paniello et al., 2012; Kato et al., 2015; Wang & Jacobsen 2017; Kato & Moynier 2017; Day et al., 2017a; 2020a). Better constraints over time and to what extent the Moon lost volatile elements will help to improve our understanding of the accretion of the Moon following the Giant Impact, as well as the accretion of other terrestrial planetary bodies. However, it is imperative to obtain rigorous constraints on the behavior of these stable isotopes in terrestrial systems, including in controlled experimental settings, before their isotopic systematics can be used to better understand complicated processes as associated with planetary scale volatile loss.

Although Ga has two stable isotopes, <sup>69</sup>Ga (60.1%) and <sup>71</sup>Ga (39.9%), few studies have investigated Ga isotope systematics in magmatic systems or in environments dominated by high temperature evaporation processes. Kato et al., (2017) analysed Ga in terrestrial igneous rocks, including the Kileauea Iki lava lake, and found the isotopic composition of Ga (δ<sup>71</sup>Ga) to be invariant at 0.00 ± 0.06‰. The same research group analysed Ga isotopes in a set of lunar basalts and lunar crustal rocks, finding a range of δ<sup>71</sup>Ga values between -0.47 and 0.85‰, with an average value for lunar basalts of 0.18 ± 0.43‰ (Kato & Moynier 2017). If there is little to no fractionation of Ga isotopes during magmatic processes such as fractional crystallization and partial melting (Kato et al., 2017), then any isotopic fractionation of Ga in lunar rocks likely occurred during evaporative loss. Kato & Moynier (2017) suggest that Ga and Zn isotope systematics behave similarly in lunar samples, with a broadly positive covariation between δ<sup>71</sup>Ga and δ<sup>66</sup>Zn values measured in the same samples (see Fig 2 in Kato & Moynier 2017). However, these data also raise some interesting, and as yet, unanswered questions about stable isotope fractionation during evaporation. For example, it is unclear why Ga isotopes are
far less fractionated than Zn in lunar basalts and highlands crustal rocks. Furthermore, it is not
84
85 known why δ<sup>71</sup>Ga values in lunar basalts vary by an order of magnitude more than in terrestrial
86
87 is to be obtained, then additional constraints on Ga isotope behavior in volatile depleted samples
88
89 from the same samples elucidating how these systems behave together during evaporation.

90 In the current study we have produced the first Ga isotope data from volatile depleted
91 terrestrial samples. By combining these data with Zn isotope data that was previously
92 presented in Wimpenny et al., (2019) we can directly evaluate the behavior of Ga and Zn during
93 evaporation from the same samples. Our data show that Ga and Zn behave similarly during
94 simple evaporative loss from silicate melts. However, in samples with more complicated
95 heritage, such as australite tektites, the δ<sup>71</sup>Ga and δ<sup>66</sup>Zn systematics are decoupled. Using both
96 Ga and Zn isotope data from volatile depleted samples allows the processes that contribute to
97 isotopic fractionation to be deduced. This has implications for the way that Ga and Zn
98 systematics are interpreted in samples from volatile depleted bodies, including for interpreting
99 isotopic fractionation on the Moon.

100

101

## 2. Method

102

### 2.1. Samples

103 The laser levitation samples analysed in this study are the same as those analysed by
104 Wimpenny et al., (2019). We undertook a series of high temperature melting experiments using
105 a gas-mixing aerodynamic levitation laser furnace at the Institut de physique du globe de Paris
106 (IPGP). The aerodynamic levitation laser apparatus is an instrument in which spherical samples
107 are freely floated on top of a stream of gas while being heated with a CO₂ laser to maximum
108 temperatures of 2200°C under atmospheric conditions, for durations of 1-120s. Levitation
109 spheres were created from rhyolitic soil from the site of a near-surface nuclear test. The soil
110 samples were taken 0.1-0.9 km from the nuclear test ground zero, away from the path of the
111 plume (Eppich et al., 2014). The soil was crushed and homogenized prior to heating using an
112 agate pestle and mortar. The rhyolitic soil was previously characterized by Eppich et al., (2014),
113 and consists of K-feldspar, quartz, plagioclase and glassy groundmass. It shows little evidence of
114 weathering or alteration. During the heating process, samples float freely on top of a stream of
115 Ar gas. Due to heating and aerodynamic rotations during heating, the samples rapidly take on a
116 spherical shape while degassing and convectively mixing. Samples heated to >1800°C formed
117 spheres that range from 0.8 to 3 mm in diameter with varying amounts of vesicularity. Samples
118 heated to temperatures <1800 °C, commonly did not form spheres, but rather melted and
formed non-aerodynamic forms. The rapidly quenched solid was then dissolved at LLNL prior to
trace element characterization and Ga and Zn isotopic analyses.

Both nuclear fallout glass and tektites form from rapid heating at extreme temperatures > 2000°C (Eby et al., 2015; Macris et al., 2019), broadly analogous to temperature-time parameters in the laser levitation experiments. We analysed the Ga isotope ratio in five fallout melt glass samples collected within 600 m of ground zero of a near-surface nuclear test that has been the subject of numerous previous studies at Lawrence Livermore National Laboratory (LLNL) (Eppich et al., 2014; Lewis et al., 2015; Weisz et al., 2017; 2018; Wimpenny et al., 2019). We also analysed Ga in four australite tektites that were previously analysed by Wimpenny et al., (2019) for Zn and major and trace elements. The tektite samples were mm in scale, with a glassy luster and sub-spheroidal in form. Finally, we performed new Ga and Zn analyses on a separate set of six australite tektites from the same strewn field. All samples were dissolved following techniques described in Wimpenny et al., (2019).

In the following we briefly describe the trace element, Zn and Ga isotope techniques used in
this study. Full details of the trace element and Zn isotope techniques are given in Wimpenny et
al., (2019). The newly developed Ga purification procedure and analytical routine at LLNL is
described here in full.

136

137 **2.2. Elemental analyses**

All sample processing and isotopic purification was performed in a class-1000 clean laboratory
at LLNL. All acids are Baseline© grade, purchased from Seastar Chemicals Inc™. Acids were
diluted using high resistivity water (18.2 MΩ) provided by Milli-Q water purification systems.
Trace element analyses were performed using the Thermo Scientific Element XR inductively
coupled plasma mass spectrometer (ICP-MS) at LLNL. All samples were digested using ultra-
pure concentrated acids (HF-HNO₃-HCl) and aliquots of stock solutions were weighed, dried
down in Teflon beakers, and diluted in an internal standard solution consisting of a 1 ppb
mixture of ¹¹⁵In, ¹⁸⁷Re, and ²⁰⁹Bi in 2% HNO₃ + 0.005M HF. Concentrations of Al, Ga and Zn were
calculated by normalizing background and drift corrected intensities to a calibration curve
prepared on the day of the analytical run. This calibration curve was created using a set of serial
dilutions of concentration standards obtained from Inorganic Ventures. To determine
measurement accuracy and reproducibility the USGS rock standard BCR-2 was analyzed along
with the sample batch. Based on the accuracy of this standard, typical uncertainties for trace
element analyses are better than 5%.

152
2.3. Zinc isotope analyses

In brief, samples were spiked using a ⁶⁶Zn-⁶⁸Zn double spike and purified by dissolving in 0.1M HCl and passing through a 0.25ml column filled by anion exchange resin (Bio-Rad AG1-X8, 200-400 mesh). Matrix elements including iron are eluted in this reagent and Zn is retained on the resin. The purified Zn is eluted in 3M HNO₃. This column is repeated to ensure a clean Zn cut. Zinc blanks are < 1ng compared to a minimum of 100-200ng of Zn processed for each sample.

Samples were dissolved in 2% HNO₃ and introduced to the plasma of a Thermo Neptune-Plus MC-ICP-MS at LLNL using an ESI Apex IR desolvating system. Typical sensitivity was 60 V of ⁶⁴Zn per ppm using a high sensitivity 'Jet' sample cone and regular 'H' skimmer cone. Samples were analysed at 50-100 ppb, with ⁶⁴Zn intensities of 3-6 V. A typical analytical routine involved multiple analyses of the IRMM-3702 standard throughout the session, followed by the QC standards (BCR-2 and BHVO-2) and unknowns.

Final data reduction followed the iterative technique developed for molybdenum isotopes in Siebert et al. (2001) where only a spiked sample is necessary to calculate the mass dependent fractionation. Although only mass dependent fractionation can be characterized using this technique, the evaporative fractionation of zinc is a mass dependent process (e.g. Moynier et al., 2009) so it is unnecessary to consider other effects.

Final δ⁶⁶Zn values are the ⁶⁶Zn/⁶⁴Zn ratio of the unknown relative to the average ⁶⁶Zn/⁶⁴Zn ratio of the bracketing standards using the following equation:

$$ \delta^{66}\text{Zn} = \left[ \left( \frac{\frac{^{66}\text{Zn}}{^{64}\text{Zn}_{\text{Sample}}}}{\frac{^{66}\text{Zn}}{^{64}\text{Zn}_{\text{IRMM 3702}}}} \right) - 1 \right] \times 1000 \qquad \text{Equation 1} $$

Zinc isotopic compositions were calculated relative to the IRMM-3702 standard. To make final data consistent with previously published work these values were then converted to the JMC-Lyon scale (δ⁶⁶ZnJMC-Lyon = δ⁶⁶ZnIRMM 3702 + 0.3‰, Cloquet et al., 2008, Petit et al., 2008, Sossi et al., 2015). Based on this method, our long-term precision can be assessed using a combination of data for BCR-2 and BHVO-2. On average, BHVO-2 had a δ⁶⁶Zn value of 0.26 ± 0.04‰ (N = 9, 2SD) and BCR-2 had a δ⁶⁶Zn value of 0.25 ± 0.04‰ (N = 14, 2SD). These values are consistent with previously published Zn isotope ratios for both standards (e.g., Bigalke et al., 2010; Chen et al., 2013; Sossi et al., 2015,). Generally, external reproducibility was ~2 times larger than the internal precision of the technique, with typical uncertainties for δ⁶⁶Zn values on individual measurements of 0.01 to 0.03‰.

183
184 **2.4. Gallium isotopic analyses**
185
2.4.1. Gallium purification
186
Gallium has two isotopes so mass bias effects cannot be corrected using a double spike.
187
Instead, isotope ratios must be bracketed to a standard, which requires high purity Ga with
188 ~100% yield through the separation process. Here, we adapted two previously published
189 procedures (Kato et al., 2017; Yuan et al., 2017) to achieve a high purity Ga fraction (see Table 1
190 for details). Using this technique we were able to analyse Ga samples down to a mass of 200ng;
191 previous studies having required at least 1000ng (Kato et al., 2017; Yuan et al., 2017). First,
192 matrix elements are removed from the sample by eluting in 6M HCl through 1.5ml of Bio-Rad
193 AG1-X8 (200-400 mesh) anion exchange resin. In strong HCl most matrix elements will pass
194 through the resin but Ga is retained. Iron is also retained on the resin at this stage, along with
195 more minor amounts of Cu, Zn and Mo. The Ga fraction is collected by eluting with 0.5M HCl.
196 The next step is to separate Ga from Fe. Here, we have adopted the technique of Yuan et al.,
197 (2017), in which Eichrom LN-spec resin is used to separate Fe from Ga. Samples are dissolved in
198 1M HCl and passed through between 1-2mL of LN-spec resin. Iron is retained on the resin and
199 Ga is eluted. The resin volume is dependent on the amount of Fe that is loaded; testing with
200 artificially doped samples showed that 5 mg of Fe could be retained by 1mL of resin. If more Fe
201 is loaded, then bleed through will occur and a second LN-spec resin step is necessary. Columns
202 1 and 2 are stick-, non-stick techniques. Therefore, it is relatively straightforward to achieve the
203 100% yield through this separation procedure. All resin from columns 1 and 2 was discarded
204 after each use.

205
Once most of the iron is removed using LN-spec resin the purified Ga still contains residual
206
levels of copper, zinc and iron. The third step is to remove these residual matrix elements using
207
Bio-Rad AG50-X12 (200-400 mesh) cation resin, similar to the technique used in Kato et al.,
208
(2017). Here, samples are dissolved in 2M HCl and eluted through 1mL of resin in glass Bio-Rad
209
Econo-columns. Any minor addition of Ba from the glass is removed in the final cleanup
210
column. The position of the Ga cut is precisely calibrated to ensure that the final Ga fraction is
211
matrix free, and that the entire cut is collected (see Fig. 1). Final Fe/Ga ratios are < 0.15. The
212
resin is retained and cleaned thoroughly using repeated 6M HCl, H₂O, 8M HNO₃ steps between
213
uses. After this, the purified Ga is dried down. However, residual organic material from the
214
resin remains in the sample, even after repeated dry downs with aqua regia. Because of the
215
sensitivity of the Ga isotope measurement to matrix effects, the purified Ga is passed through a
216
final anion clean up column, similar to column-1 but with a reduced resin volume (~0.2 mL). The
217
final Ga fraction is dried down and redissolved in 2% HNO₃ ready for analysis. Total Ga
218
procedural blanks, including digestion, are < 0.05 ng. Where Zn and Ga are purified from the
219
same sample aliquot the Zn is first separated from matrix elements in 0.1M HCl on anion resin.
The matrix aliquot contains the Ga, and this is processed through the full separation procedure outlined above.

### 2.4.2. Gallium isotopic analysis

Gallium isotope ratio measurements were made using the Thermo Scientific Neptune-plus MC-ICP-MS at LLNL. To reduce the potential for sample matrix effects to introduce variable mass bias, the sample is introduced in wet plasma mode. Using the high sensitivity 'Jet' sample cone and standard 'H' skimmer cone ~2.5V of ⁶⁹Ga for a 40 ng/g sample was obtained. Although the sensitivity is relatively low compared to dry plasma techniques, the fact that Ga has two relatively abundant isotopes (60.1 and 39.9%) means ⁷¹Ga was still analysed at >1.5V. The detector array was set up to monitor ⁶³Cu⁺, ⁶⁵Cu⁺, ⁶⁶Zn⁺, ⁶⁷Zn⁺, ¹³⁷Ba²⁺ (mass 68.5), ⁶⁹Ga⁺ and ⁷¹Ga⁺ on faraday cups with 10¹¹-ohm resistors. The most critical species to monitor was doubly charged barium, as doubly charged ions from the major isotope of barium (¹³⁸Ba, 71.7%) interfere with ⁶⁹Ga. Barium-137 is a relatively low abundance isotope (11.2%), hence any Ba correction applied to ¹³⁸Ba requires the multiplication of signal at mass 68.5 by 6.4 (71.7%/11.2%). This method of correction is highly problematic. The background signal of 68.5 varied in blanks and pure Ga standards between 50 and 100 µV. Hence, false positives on the 68.5 peak were observed as high as 50 µV after blank subtraction. Although such false positive signals are rare, the false subtraction of a 50 µV signal on a 2.5 V peak will shift the ⁷¹Ga/⁶⁹Ga isotope ratio by ~0.2‰. This is unacceptable considering that the average δ⁷¹Ga value of the BSE is estimated at 0.00 ± 0.06‰. To resolve this problem, all samples that yielded a measurable 68.5 signal, outside typical background values, were re-purified. No trace of barium in the final samples was observed after adopting the final cleanup column (which also removes barium).

To correct for instrumental mass bias effects the sample was bracketed to a Ga standard. For Ga isotope data to be easily comparable with literature data it is helpful if this bracketing standard is an internationally available reference material. Previous studies by Kato et al., (2017) and Kato & Moynier (2017) used an in-house Ga standard at IPGP, whereas studies by Zhang et al., (2016), Yuan et al. (2017) and Feng et al., (2019) had used the NIST Ga isotope standard SRM 994. Due to the wide availability of the NIST standard, all data are bracketed using SRM 994. To monitor instrumental performance, we analysed an in-house Ga standard (from Inorganic Ventures) multiple times during each analytical session (Fig. 2a). This has a δ⁷¹Ga value of 1.24 ± 0.07‰ (n = 51). Whereas NIST 994 is designated as an isotopic standard, it is certified only to ± 0.25‰ and is shipped in metallic form, in which isotopic heterogeneities can be preserved (e.g. Galy et al., 2003). Multiple analyses of basaltic standards BCR-2 and BIR-1a resulted in δ⁷¹Ga values of +0.50‰ relative to NIST 994 (Fig. 2). In comparison, Yuan et al., (2017) obtained basaltic values of +0.10‰ relative to SRM 994 and Zhang et al., (2016) and
Feng et al., (2019) obtained δ⁷¹Ga values of +0.75‰. Note that Zhang et al., (2017) and Feng et al., (2019) used the same batch of SRM 994 from the same institution. Yuan et al., (2017) validated their data by performing a standard addition experiment by mixing SRM 994 with an in-house pure Ga solution at different proportions and processing these through their chemical purification procedure. The measured isotopic composition of the mixtures should ideally yield a straight line when plotted as a function of their mixing ratios. We performed a similar test, as shown in Fig. 3, in which we mixed SRM 994 and BCR-2 at different proportions, purified the sample through the ion-exchange procedure described above, and analysed the final Ga. This is a far better test, because it also incorporates the effect of matrix elements on the chemical separation and final isotopic analyses. As shown, the data lie on a straight line that lies within uncertainty of the BCR-2 endmember at +0.50‰ and the SRM 994 endmember at 0‰. To validate our technique further we analysed the USGS rock standards AGV-2 (andesite), RGM-2 (rhyolite), SDC-1 (mica schist) and GSP-1 (granodiorite), together with a Syenite powder from the Canadian Geological Survey. Despite variations in chemical composition and sample matrix, all of the final δ⁷¹Ga values of the rock standards were within uncertainty of 0.50 ‰. Based on this, our technique appears to be robust, suggesting that SRM 994 is isotopically heterogeneous at the ‰ level. It is therefore not a good bracketing standard for Ga isotopic analyses for future investigations. Advancing Ga isotope geochemistry will require the choice of a suitable pure Ga standard that is isotopically homogenous and widely available to the scientific community. We recommend that this standard be distributed in solution form, rather than a metal to reduce the likelihood of isotopic heterogeneity. In the meantime, we have normalized all of our final data to the δ⁷¹Ga value of terrestrial basalts BCR-2, BHVO-2 and BIR-1a, which are within uncertainty of 0.50‰ relative to our aliquot of SRM 994 (Fig. 2a). This allows our data to be directly comparable to that of Kato et al., (2017) and Kato & Moynier (2017). Thus, our final calculation of δ⁷¹Ga involves the following two equations:

$$ \delta^{71}Ga_{NIST 994} = \left[ \left( \frac{\frac{71Ga}{69Ga_{Sample}}}{\frac{71Ga}{69Ga_{NIST 994}}} \right) - 1 \right] \times 1000 \qquad \text{Equation 2} $$

$$ \delta^{71}Ga_{IPGP} = \delta^{71}Ga_{NIST 994} - 0.50\text{‰} \qquad \text{Equation 3} $$

### 3. Results

Concentrations and isotopic compositions of Ga and Zn, as well as Al abundances, obtained on the samples is presented in Table 2. All δ⁷¹Ga values are calculated relative to the δ⁷¹Ga value of +0.50‰ determined relative to SRM 994 (see Eq. 3), which is a reasonable proxy for the BSE
(see also Kato et al., 2017). Major element concentrations for the experimental samples and previously analysed australites are presented in Wimpenny et al., (2019).

## 3.1. Heating experiments

The rhyolitic soil used in these experiments derives from the vicinity of a near surface nuclear test (see Section 3.2.). It contains 29 ppm Ga, which is relatively enriched compared to typical upper continental crust (UCC) at 17.5 ppm (Rudnick & Gao 2003). The rhyolitic soil was heated to temperatures between 1600-2200 °C for between 1-120 s. Significant changes in major and trace element concentration correlate with temperature and time of heating. Refractory elements, such as Ca and Al, typically increased in concentration with heating temperature and duration whereas volatile elements, including K, Na, Pb, Tl, and Zn, decreased (Fig. 4 and 5, Wimpenny et al., 2019). Gallium concentration in the laser spheres ranges between 30 ppm and 19 ppm, progressively decreasing with temperature of the experiment, so that samples heated to 2200 °C have the lowest Ga concentrations. In general, the Ga concentration in these spheres correlates with the abundance of the volatile element Zn, and anticorrelates with abundance of the refractory element Al.

The $\delta^{71}$Ga value of the unheated rhyolitic soil is -0.03 ± 0.06 ‰, within error of the BSE, and consistent with the relative isotopic homogeneity of Ga in crustal materials (Kato et al., 2017). In comparison, the laser spheres have $\delta^{71}$Ga values that range from -0.01 to +0.80 ‰, i.e. always enriched in isotopically heavy Ga relative to the starting soil. Furthermore, the samples that are most fractionated from the starting soil also have the lowest Ga concentrations (Fig. 4). The Zn and Ga isotope systematics of the laser spheres are broadly similar. In both cases the laser sphere heated to 2200°C for 60s (rhyolite soil-15) contains the most fractionated $\delta^{71}$Ga and $\delta^{66}$Zn values of +0.80 and +1.49‰, respectively.

## 3.2. Fallout melt glass

Five fallout melt glass samples were recovered from a near surface nuclear test that had interacted with the same rhyolitic soil homogenized and used in the laser heating experiments. Consequently, the Ga concentration in the starting soil is the same as used in the experimental work (29 ppm). Relative to their rhyolitic soil precursor, the fallout samples have roughly similar chemical compositions, but are depleted in the volatile elements Na, K, Zn, Tl, and Pb, and enriched in the refractory elements Ca and Al (Wimpenny et al., 2019). Gallium concentration in the fallout samples range from 12.8 to 24.2 ppm and are depleted compared to the rhyolitic
soil. The δ<sup>71</sup>Ga of the fallout samples ranges from 0.05 to 0.38‰, with a negative correlation between Ga concentration and δ<sup>71</sup>Ga, as shown in Fig. 4.

323

324 **3.3. Tektites**

Gallium concentrations in the australites range from 5.9 to 8.7 ppm and are relatively low compared to UCC values of ~17 ppm (Rudnick & Gao 2003). The Ga concentration positively correlates with most major element concentrations, which in turn, negatively correlate with silica content. Negative correlations between silica and most major oxides is a well-known phenomenon in australites (e.g. Koeberl 1992). The δ<sup>71</sup>Ga value of these samples range from 0.04 to 0.38 ‰, with a negative correlation between δ<sup>71</sup>Ga and [Ga] as shown in Fig 4.

331

332 **4. Discussion**

333 **4.1. Gallium isotope systematics in laser heating experiments**

Major and trace element analyses of the experimentally heated rhyolitic soil demonstrates that they progressively lost more volatile elements, e.g. K, Rb, Pb, Tl and Zn, with increasing temperature and time (see trace element data in Wimpenny et al., 2019). The proportion of elemental loss cannot be directly calculated because the initial sample mass used in the experiment is not well constrained. Instead, it is only possible to calculate element loss after normalizing to an assumed refractory index element, which is likely to be conserved during heating. Here we have chosen Al as the index element due to its relatively high 50% condensation temperature (1403 °C, Lodders 2003), high abundance, and extreme apparent enrichments in the laser spheres heated to 2200°C (Wimpenny et al., 2019). The use of Al as a normalizing element is not ideal because some non-negligible loss of Al during heating cannot be ruled out.

There are relationships between the loss of more volatile elements, such as Pb, Tl and Zn, with heating temperature and duration when the data are normalized to Al (Wimpenny et al., 2019). The laser spheres demonstrate between 0 – 48% loss of the initial Ga, with the most Ga lost in sample rhyolite-15, which had the most extreme T-t conditions (2200°C, 60s). The relationship between the percentage Ga loss and δ<sup>71</sup>Ga for each experimental sample is presented in Fig. 5a, alongside the corresponding data for Zn from Wimpenny et al., (2019). Note that there are two outlying Ga samples (rhyolite soil-8 and -10) that have anomalously heavy δ<sup>71</sup>Ga values. We interpret these offsets as mass bias effects during mass spectrometry. These samples contained the lowest mass of Ga processed through ion exchange chromatography (~50 ng) and were run

346

347

348

349

350

351

352

353
with minimal dilution, which enhances matrix differences between sample and standard. As a consequence, we have excluded them from further discussion of the Ga fractionation factor.

There is a clear relationship between the proportion of Ga loss and δ<sup>71</sup>Ga value so that samples with higher δ<sup>71</sup>Ga values have lost more Ga. This trend matches that observed for Zn isotopes in the same samples and is consistent with the fact that both elements are moderately volatile and that both isotope systems undergo kinetic fractionation during evaporation.

The fractionation factor (α) associated with evaporative loss can be calculated for Ga based on the following relationship:

$$R = R_0 \times f^{\alpha-1} \qquad \text{Equation 4}$$

Where R is the Ga isotope ratio of the volatile depleted experimental sample, R₀ is the initial Ga isotope ratio in the bulk soil and f is the fraction of Ga that remains after evaporation. As shown in Fig. 6, if the natural log of R/R₀ is plotted versus the natural log of the remaining Ga fraction, the data define a straight line, the slope of which can be used to calculate the Ga isotopic fractionation factor. The α calculated from the slope in Fig. 6 is 0.99891 ± 0.00024 (2 SD), which is within uncertainty of the fractionation factor calculated for Zn from the same samples (α = 0.99879 ± 0.00013, Wimpenny et al., 2019).

The theoretical fractionation factor for evaporation of an element, in this case Ga, in a vacuum is estimated based on the following equation:

$$\alpha = \sqrt{\frac{m^{69}Ga}{m^{71}Ga}} \qquad \text{Equation 5}$$

Where m is the mass of the isotope or evaporating species.

Thus, the theoretical α is 0.9858 for evaporation of elemental Ga and 0.9847 for evaporation of elemental Zn. However, the α value is, to a certain extent, dependent on the speciation of the evaporating Ga. In addition to evaporating as a metal, gallium forms sulphide and chloride gases in igneous systems (Renggli et al., 2017). Based on their most abundant isotopes (³⁵Cl = 75.76%, ³²S = 94.99%), the theoretical α for evaporation of GaCl₃ and GaS are 0.994 and 0.990 respectively. In both cases the expected fractionation of Ga is significantly lower than our experimental observations. The suppression of the Ga fractionation observed in the experiments is consistent with the behavior of Zn (Wimpenny et al., 2019), as well as previous experimental studies investigating the kinetic fractionation of Mg, Si, and O during evaporation of olivine and refractory inclusions in chondritic meteorites (e.g. Davis et al., 1990; Wang et al., 1999; Grossman et al., 2000, 2002; Richter et al., 2002, 2007, 2011; Knight et al., 2009; Mendybaev et al., 2017). The likely cause of this suppression is the environment in which
386 evaporation occurred because it is well established that the theoretical kinetic fractionation 387 factor is only attained experimentally at high vacuum (Tsuchiyama et al., 1999; Richter et al., 388 2002). At higher pressures the effective fractionation factor α' is approximated by the following 389 equation from Richter et al., (2002):

$$ (\alpha' - 1) \approx (\alpha - 1) \left( 1 - \frac{P_i}{P_{i,sat}} \right) \qquad \text{Equation 6} $$

391 Where Pᵢ is the pressure of gas species i at the evaporating surface and Pᵢ,sat is the equilibrium 392 vapor pressure of i (Richter et al., 2002, Young et al., 2019). As both pressures increase the 393 value of Pᵢ/Pᵢ,sat approaches 1, which results in an effective fractionation factor (α') of 1 (i.e. no 394 isotopic fractionation). This is caused by recondensation of evaporated material back onto the 395 sample surface, resulting in lower isotopic fractionation until the system reaches equilibrium 396 (Richter et al., 2002). This is discussed further for Zn isotope systematics in Wimpenny et al., 397 (2019).

398 The similarity between experimentally determined α values for Ga and Zn is not surprising given 399 that the relative mass difference between Ga and Zn isotopes is similar. Because there is a 400 slightly larger mass difference between Zn isotopes than Ga isotopes, a given fraction of 401 evaporative loss is expected to induce a slightly greater shift in δ⁶⁶Zn. Therefore, in δ⁷¹Ga-δ⁶⁶Zn 402 space (with δ⁶⁶Zn on the x-axis) samples that have lost equal proportions of Ga and Zn should 403 lie on a line of slope ~0.93, which, for simplicity, is identified here as the 1:1 line (Fig. 7). 404 However, it is important to consider that Zn is more volatile than Ga (50% Tc of 453 °C and 695 405 °C respectively, Lodders 2003) and a greater fraction of Zn loss will impart a greater isotopic 406 effect in the residual glass. Because the rhyolitic soil lost between 10 and 20% more Zn than Ga 407 the experimental data are predicted to lie to the right of the 1:1 line. As shown in Fig. 7 our 408 isotopic data are consistent with this prediction.

409 These data illustrate that the direct comparison of moderately volatile isotope data can provide 410 a useful method to evaluate the processes that cause isotopic fractionation and the relative 411 extent of element depletion in volatile depleted samples. In this case the experimental system 412 is relatively simple using a homogenous rhyolitic soil that exhibits little evidence of weathering 413 or alteration that might impart geochemical heterogeneity or induce melting at different 414 liquidus temperatures (Eppich et al., 2014). In more complicated samples with source 415 composition and/or isotopic heterogeneities or mineralogical differences, the δ⁷¹Ga- δ⁶⁶Zn 416 systematics may be more difficult to interpret.

417
418
**4.2. Gallium isotope systematics in fallout melt glass**
During a nuclear detonation initial temperatures and pressures exceed 8000°C and 8GPa (e.g. Glasstone & Dolan 1977; Hermes & Strickfaden, 2005; Eby et al., 2015, Roberts et al., 2019). In near surface tests the fireball entrains local materials, such as soil, which undergoes rapid heating to temperatures between 1200-2000°C (Cassata et al., 2014; Eby et al., 2015; Macris et al., 2019), undergoes melting and assimilation of melt droplets (e.g. Weisz et al., 2017; 2018; Lewis et al., 2017), and loses a significant proportion of its volatiles. Based on the thermal history of entrained environmental material from near surface blasts it is intuitive that fallout melt glass has lower concentrations of relatively volatile elements including Ga, Zn, Pb and Tl, compared to its parent soil (Wimpenny et al., 2019). After normalization to the concentration of Al the five fallout samples studied here lost between 6 and 53% of the initial Ga. The relationship between Ga loss and Ga isotope fractionation is plotted in Fig. 5b. In general, there is a positive relationship between the proportion of Ga loss and fractionation of the Ga isotope ratio with increasing Ga depletion corresponding to increasing enrichment in 71Ga. This behavior is consistent with kinetic fractionation of Ga isotopes during evaporation and matches the behavior of Zn isotopes in fallout melt glass (Day et al., 2017b; 2020b; Wimpenny et al., 2019). A plot of the natural log of R/R₀ versus the natural log of the remaining Ga fraction demonstrates that the data are relatively scattered and do not define a well constrained regression. However, as shown in Fig. 5b the Ga isotope data are generally consistent with a relatively suppressed fractionation factor of ~ 0.9997, similar to Zn isotope data from the same samples (Wimpenny et al., 2019) and Trinitite glasses (Day et al., 2017b). It is possible that the fallout melt glass is affected by anthropogenic input of Ga or Zn with distinctive stable isotope ratios (e.g. John et al., 2007). However, in this case the degree of isotopic fractionation is expected to increase with Ga and Zn concentration, which is opposite the trends observed in our data (Fig 4a and b).

Following the discussion in Section 4.1, and Section 4.3.2 of Wimpenny et al., (2019), the relative suppression of α in fallout melt glass, compared to the laser heating experiments, is interpreted to reflect differences between the physical environments of evaporation. While the vapor pressure within the debris cloud from a nuclear test is not easily constrained, it is reasonable to assume that it is higher than ambient (see also Day et al., 2020b). In contrast, the laser heating experiments were performed at atmospheric pressure, and samples were levitated on a stream of argon gas, which would have transported evaporated species away from the glass surface, reducing the amount of recondensation. Thus, it is likely that the Pᵢ/Pᵢ,sat was closer to 1 in the debris cloud suppressing isotopic fractionation associated with evaporation in the nuclear fallout glass samples (e.g. Richter et al., 2002, 2007).

If Ga and Zn isotope systematics are directly compared by plotting them in δ⁷¹Ga-δ⁶⁶Zn space (Fig. 7) the fallout data form an array that is similar to the experimental laser fusion samples; i.e. a straightforward correlation between δ⁷¹Ga and δ⁶⁶Zn that is broadly consistent with
isotopic fractionation by evaporative loss is observed. This is expected given that the fallout is
dominantly composed of rhyolitic soil of the same composition as that used in the laser
levitation experiments. However, this also suggests that any differences in physical conditions
(i.e. P, T, fO₂) between the laser heating experiments and fallout formation in the fireball do not
significantly change Zn and Ga isotope systematics.

461

462 **4.3. Gallium isotopic systematics in australite tektites**

Estimates for the proportion of Ga loss from australite tektites are controlled by the estimates
for the bulk chemical composition of the target rock. Unfortunately, the location of the
Australite impact crater is still debated and locations including eastern Thailand, the Gulf of
Tonkin, and the Chinese Loess Plateau have all been proposed (e.g. Schnetzler 1992; Glass &
Koeberl 2006; Mizera et al., 2016). Although there is a myriad of potential lithologies, analyses
of trace element ratios, Rb-Sr isotope systematics, and ¹⁰Be concentrations have established
that the parent rock was probably a weathered sedimentary material (e.g. Koeberl 1992; Ma et
al., 2004; Mizera et al., 2016). An estimate of the composition of the target rocks is obtained by
taking the average composition of Muong-Nong type tektites from the Australasian strewn
field. These are more primitive, volatile rich tektites that are hypothesized to have formed co-
genetically with the splash-form australites (Schnetzler 1992; Koeberl 1992). The average
concentration of Ga in these Muong-Nong type tektites is 23 ± 11 ppm (Koeberl 1992), which is
within uncertainty of the average UCC of 17.5 ± 1.4 ppm (Rudnick & Gao 2003). Based on a
Muong-Nong type precursor, evaporation of between 55-65% of the initial Ga is calculated. In
comparison, between 62-89% of Zn is calculated to have been evaporated assuming the
protolith of the tektites had a concentration of 90 ppm as measured in Muong-Nong australites
(Koeberl et al., 1992). This means that approximately 20-30% more of the initial Zn was
volatilized than Ga, which is broadly consistent with the behavior of these elements during
experimental heating (section 4.2). However, no clear relationship exists between the
proportion of Ga and Zn loss in the tektite samples, which is surprising given their relative
volatilities and similar behavior during experimental heating.

463

464

465

466

467

468

469

470

471

472

473

474

475

476

477

478

479

480

481

482

483

The Ga isotope systematics in australites do not form a coherent array and so cannot be used to
define a precise fractionation factor. The lack of spread in either the proportion of Ga loss or
δ⁷¹Ga values means that any attempt to calculate α results in a two-point regression with the
origin. Thus, a meaningful α cannot be calculated from Ga isotopic systematics in australite
tektites, consistent with previously published data for Zn (Wimpenny et al., 2019).

484

485

486

487

488

489

490

Australites and fallout melt glass are hypothesized to have formed in broadly similar
environments with rapid, short duration heating to temperatures > 2000°C (Eby et al., 2015;
Macris et al., 2019), pressures above 8GPa (French et al., 1998; Eby et al., 2015), and reducing conditions, as evidenced from low Fe³⁺/Fe²⁺ ratios in both types of sample (Glass 1987; Belkin & Horton 2009). However, when plotted together, the δ⁷¹Ga-δ⁶⁶Zn systematics of australite tektites are offset to the right of the fallout melt glass and the experimental array (Fig. 7), due to the greater extent of Zn isotopic fractionation. Furthermore, the δ⁷¹Ga-δ⁶⁶Zn systematics of australite tektites do not positively correlate, as is expected if simple kinetic fractionation took place and as is observed for both the laser spheres and nuclear fallout glass (Fig. 7). Taken together, this suggests that the offset and scatter in the δ⁷¹Ga-δ⁶⁶Zn systematics of australites is not caused simply by a difference in the physical conditions of evaporation.

A more likely explanation is that the δ⁷¹Ga-δ⁶⁶Zn systematics of tektites reflect isotopic fractionation of the tektite precursor lithology. Unlike the rhyolitic soil used in the laser heating experiments, which is composed of primary silicate minerals and is relatively free of alteration (Eppich et al., 2014), the parent rock to the australites is likely to be weathered sedimentary material including both silicates and carbonates (e.g. Koeberl 1992; Ma et al., 2004; Mizera et al., 2016). The presence of dark and light banding in Muong Nong type australites implies that at least two parent lithologies with different silica contents were incorporated into the debris cloud during tektite formation (Koeberl 1992, Son & Koeberl 2005). This is supported by the fact that major element oxide abundances negatively correlate with the concentration of Si in most tektite groups (Taylor et al., 1962; Koeberl 1988, Son & Koeberl 2005; Amare & Koeberl 2006; Schwarz et al., 2016). The target sediment is hypothesized to be a mixture of quartz and shale-like materials, with variable addition of carbonate (Howard 2011 and references within).

The incorporation of a such a heterogeneous mix of weathered material is potentially an important control over the isotopic systematics of tektites. Weathering processes are well known to impart isotopic heterogeneities in the composition of many stable isotope systems, including Zn in igneous rocks, sediments, and surface waters (Pokrovsky et al., 2005; Fernandez & Borrok 2009; Guinoiseau et al., 2014; Lv et al., 2016; Mavromatis et al., 2019). For example, Zn that is adsorbed onto kaolinite is between 0.2-0.5‰ heavier than the corresponding fluid phase (Guinoiseau et al., 2014), whereas the fractionation is slightly larger during adsorption onto calcite (0.4-0.7‰, Dong & Wasylenki, 2016; Mavromatis et al., 2019). Although there is little information about Ga isotope fractionation during weathering processes, recent work by Yuan et al., (2018) demonstrated that Ga taken onto calcite and goethite is approximately 1.1 and 0.6‰ lighter than Ga in the fluid phase, respectively. Thus, weathered sediment is likely to contain isotopically heterogeneous δ⁷¹Ga-δ⁶⁶Zn systematics with the available data suggesting that weathered materials will form a negative array in δ⁷¹Ga-δ⁶⁶Zn space. A plausible explanation for the scattered δ⁷¹Ga-δ⁶⁶Zn systematics of australite tektites is that it reflects a combination of isotopic heterogeneity of the source material together with kinetic fractionation associated with variable degrees of evaporative loss. Note that post-formation weathering is
unlikely to affect the Zn and Ga isotope composition of australite tektites, which are well
preserved with little evidence of chemical perturbation (Taylor & Sachs 1964). The importance
of the precursor lithology in contributing to variations in δ<sup>71</sup>Ga and δ<sup>66</sup>Zn values is consistent
with previous studies of Li (Magna et al., 2006) and B isotopes (Chaussidon & Koeberl 1995) in
tektites.

The fact that the tektite data is offset from the experimental array is consistent with the compositional heterogeneities in the australite precursor, as well as the likely presence of secondary silicates and carbonate minerals in the precursor rocks. It is well known that carbonate-rich, hydrated melts have lower liquidus temperatures (500-600 °C, Osinski & Spray 2001) than silicate melts, which have liquidus temperatures of ~1200 °C (e.g. McKenzie 1985; Jones et al. 2013). The melting and partial evaporation of a more reactive phase that contained less silica, but a relatively high Zn/Ga ratio would result in greater extent of Zn loss and associated kinetic fractionation. Of course, this scenario requires only partial loss of the more reactive phase because quantitative evaporative loss would not result in isotopic fractionation.

542

543

5. Implications for understanding δ<sup>71</sup>Ga- δ<sup>66</sup>Zn systematics in the Moon

Previous studies of MVE isotope systems in mare basalts, including Zn (Paniello et al., 2012; Kato et al., 2015), Cu (Moynier et al., 2006; Herzog et al., 2009), K (Wang & Jacobsen 2016; Tian et al., 2020), and Rb (Pringle & Moynier 2017; Nie & Dauphas 2019) have found that these elements are isotopically heavier than the BSE, consistent with the Moon's relative depletion in volatile elements and compounds (e.g. Wänke et al., 1973; Ringwood & Kesson, 1977; Newsom & Taylor 1989; Albarede 2009). This is significant because mare basalts are often used as a proxy for the BSM (e.g. Nie & Dauphas 2019; Tian et al., 2020). Although some of the isotopic differences between the Moon and Earth could be attributed to volatile loss caused by the Giant Impact event (e.g. Paniello et al., 2012; Wang & Jacobsen 2016), isotopic fractionation of MVEs can also be caused by post-accretion processes. Modification of MVE isotope compositions resulting from post-accretion processing is manifest by isotopic heterogeneities in both mare basalts and rocks from the lunar crust (Kato et al., 2015; Day et al., 2017a; 2020a). Such isotopic heterogeneities are particularly significant for Ga; the average δ<sup>71</sup>Ga value of mare basalts (+0.18 ± 0.43‰, Kato & Moynier 2017) currently cannot be resolved from the BSE (0.00 ± 0.06‰, Kato et al., 2017). It's important to constrain the processes that cause MVE isotopic heterogeneities in order to make robust estimates of the isotopic composition of the BSM.

544

545

546

547

548

549

550

551

552

553

554

555

556

557

558

559

560

561

562

The experimental data presented here, together with observations from nuclear fallout glass, predicts that evaporative loss of MVEs from a silicate melt will cause fractionation of δ<sup>71</sup>Ga that
563 positively correlates with δ⁶⁶Zn (Fig. 7). However, as shown in Fig. 8 the published δ⁷¹Ga and
564 δ⁶⁶Zn values of mare basalts are scattered (Kato et al., 2015; Kato & Moynier 2017), suggesting
565 that the two isotopic systems are decoupled. This implies that simple evaporative fractionation,
566 either during degassing of the lunar magma ocean or accompanying eruptive degassing, cannot
567 be the sole fractionating mechanism. Similarly, extensive degassing of the ur-KREEP reservoir by
568 crust breaching impact (Barnes et al., 2016) may result in isotopic fractionation of Zn (Day et al.,
569 2020a) and/or Ga, but variable mixing of this fractionated signal back into the mare basalt
570 source regions should, in theory, still generate correlated δ⁷¹Ga-δ⁶⁶Zn systematics.

571 A plausible mechanism that could result in decoupling of δ⁷¹Ga-δ⁶⁶Zn systematics in lunar
572 samples is the redistribution of volatiles by impacts at the Moon's surface. Evidence suggests
573 that this has a significant effect on Zn isotope systematics in rocks from the lunar crust, which
574 are often heavily altered and brecciated by impacts, and can contain volatile rich impact melt
575 glass (e.g. Shearer et al., 2014; Day et al., 2017a). Zinc concentrations in ferroan anorthosites
576 and impact melt breccias can be up to two orders of magnitude higher than the majority of
577 mare basalts (Kato et al., 2015; Day et al., 2017a). Zinc isotope ratios are also highly variable
578 with δ⁶⁶Zn values that range from -13.7 to +4‰ (Kato et al., 2015; Day et al., 2017a). To date,
579 there is little variation in either the concentration of Ga (Neal et al., 2001; Hallis et al., 2014;
580 Kato & Moynier 2017) or its isotopic composition in crustal rocks (Kato & Moynier 2017). This
581 implies that Zn may be more easily mobilized than Ga during impact reworking, likely because
582 of differences in volatility and speciation during evaporation (Lodders 2003; Renggli et al.,
583 2017).

584 Although the addition of vapor condensate can explain anomalously negative δ⁶⁶Zn values in a
585 small number of mare basalts, such as 10017 and 12018 (Day et al., 2020a), most mare basalts
586 have a relatively narrow range of δ⁶⁶Zn values between 0.8-1.9 ‰ (Herzog et al., 2009; Paniello
587 et al., 2012; Kato et al., 2015; Day et al., 2020a), indicating that condensate addition is far less
588 important. If we assume that Ga is less easily mobilized than Zn, this suggests that δ⁷¹Ga values
589 of mare basalts are unlikely to have been perturbed by condensate addition. Thus, an
590 alternative mechanism is required to explain the Ga isotopic heterogeneity shown in Fig 8. One
591 possibility is that Ga isotope systematics are sensitive to mineral-melt partitioning during
592 solidification of the lunar magma ocean (LMO). However, currently there are no constraints
593 over the extent of Ga isotope fractionation between coexisting mineral phases. To test this
594 hypothesis new studies are needed, focusing on the behaviour of Ga in minerals such as
595 plagioclase, clinopyroxene and spinel, which are important crystallization products of the LMO
596 (e.g. Snyder et al., 1992), and in which Ga is relatively compatible (Malvin & Drake 1987).

597 Ultimately, we draw two conclusions from the lunar δ⁷¹Ga-δ⁶⁶Zn data. First, isotopic
598 heterogeneities in δ⁷¹Ga and δ⁶⁶Zn in mare basalts must be generated by processes that post-
date lunar accretion. Therefore, the average isotopic composition of mare basalts may not reflect the isotopic composition of the BSM. Second, there is not a straightforward relationship between δ⁷¹Ga and δ⁶⁶Zn systematics in mare basalts, contrary to the behavior of these elements in our experiments. It therefore follows that characterizing the behavior of a single MVE isotope system in lunar rocks will not necessarily provide accurate information about the loss of other MVEs from the Moon.

## 6. Conclusions

New analyses of Ga isotope ratios in terrestrial volatile depleted materials show that Ga isotopes are sensitive indicators of volatile loss. Based on laser levitation experiments, we calculate that evaporative fractionation of Ga isotopes took place with an α value of 0.99891 ± 0.00024, which is within uncertainty of the α value calculated for Zn isotopes in the same samples. Analyses of Ga isotope ratios in nuclear fallout glass are consistent with a relatively suppressed α and are similar to results of Zn isotopic analyses. These data highlight that when simple silicate-rich materials, such as the rhyolitic soil, are subjected to rapid high temperature heating the isotopic composition of moderately volatile elements such as Zn and Ga behave very similarly.

The situation is more complicated in australites, because good constraints over the target rocks are scarce precluding confident determination of α. Furthermore, Ga and Zn isotope ratios do not positively correlate as expected if evaporative loss is the sole control over isotopic fractionation. Instead, we hypothesize that scattered δ⁷¹Ga-δ⁶⁶Zn systematics in australites reflect isotopic heterogeneity in the target rocks, coupled with variable evaporative loss.

In theory, evaporative loss of Ga and Zn from the Moon, either pre- or post-accretion should also be a relatively simple process involving evaporation from a silicate melt. Hence the Ga and Zn isotope systematics in lunar basalts are expected to positively correlate. Instead the published data are scattered and offset from the experimentally derived δ⁷¹Ga-δ⁶⁶Zn array. This suggests δ⁷¹Ga-δ⁶⁶Zn systematics were affected by processes that post-date lunar accretion. Hence, the average isotopic composition of mare basalts may not reflect the isotopic composition of the BSM.

## 629 Acknowledgements

The study benefited from the constructive comments of three anonymous reviewers and the associate editor, James Day. This work was funded by the LLNL-LDRD program under project
632 number LW-LDRD-027 and by the Office of Defense Nuclear Nonproliferation Research and
633 Development within the U.S. Department of Energy's National Nuclear Security Administration.
634 Nuclear test site samples were made available from collections at LLNL with the permissions of
635 Office of Defense Nuclear Nonproliferation Research and Development within the U.S.
636 Department of Energy. This work was performed under the auspices of the U.S. Department of
637 Energy by Lawrence Livermore National Laboratory under Contract DE-AC52-07NA27344. LLNL-
638 JRNL-798685

639 JB acknowledges the financial support of the UnivEarthS Labex program at Sorbonne Paris Cité
640 (ANR-10-LABX-0023 and ANR-11-IDEX-0005-02) and IPGP program PARI/Paris-IdF region
641 SESAME Grant no. 12015908.

642
643 References

644 Albarède, F., Albalat, E., & Lee, C. T. A. (2015). An intrinsic volatility scale relevant to the Earth and Moon
645 and the status of water in the Moon. *Meteoritics & Planetary Science*, 50(4), 568-577.

646 Amare, K., & Koeberl, C. (2006). Variation of chemical composition in Australasian tektites from different
647 localities in Vietnam. *Meteoritics & Planetary Science*, 41(1), 107-123.

648 Armytage, R. M. G., Georg, R. B., Williams, H. M., & Halliday, A. N. (2012). Silicon isotopes in lunar rocks: Implications for the Moon's formation and the early history of the Earth. *Geochimica et Cosmochimica Acta*, 77, 504-514.

651 Barnes, J. J., Tartese, R., Anand, M., Mccubbin, F. M., Neal, C. R., & Franchi, I. A. (2016). Early degassing of lunar urKREEP by crust-breaching impact (s). Earth and Planetary Science Letters, 447, 84-94.

653 Belkin, H. E., & Horton Jr, J. W. (2009). Silicate glasses and sulfide melts in the ICDP-USGS Eyreville B core, Chesapeake Bay impact structure, Virginia, USA. *The ICDP-USGS deep drilling project in the Chesapeake Bay impact structure: Results from the Eyreville core holes*, 458, 447.

656 Bigalke, M., Weyer, S., Kobza, J., & Wilcke, W. (2010). Stable Cu and Zn isotope ratios as tracers of sources and transport of Cu and Zn in contaminated soil. *Geochimica et Cosmochimica Acta*, 74(23), 6801-6813.

659 Blanchard, I., Badro, J., Siebert, J., & Ryerson, F. J. (2015). Composition of the core from gallium metal-silicate partitioning experiments. *Earth and Planetary Science Letters*, 427, 191-201.

661 Borg, L. E., Gaffney, A. M., Kruijer, T. S., Marks, N. A., Sio, C. K., & Wimpenny, J. (2019). Isotopic evidence for a young lunar magma ocean. *Earth and Planetary Science Letters*, 523, 115706.

663 Butler Jr, J., & Meyer Jr, C. (1976, April). Sulfur prevails in coatings on glass droplets-Apollo 15 green and brown glasses and Apollo 17 orange and black/devitrified/glasses. In *Lunar and Planetary Science Conference Proceedings* (Vol. 7, pp. 1561-1581).

666 Butler Jr, P. (1978). Recognition of lunar glass droplets produced directly from endogenous liquids-The evidence from S-ZN coatings. In *Lunar and Planetary Science Conference Proceedings* (Vol. 9, pp. 1459-1471).

669 Cassata, W. S., Prussin, S. G., Knight, K. B., Hutcheon, I. D., Isselhardt, B. H., & Renne, P. R. (2014). When the dust settles: stable xenon isotope constraints on the formation of nuclear fallout. *Journal of environmental radioactivity*, 137, 88-95.

672 Chaussidon, M., & Koeberl, C. (1995). Boron content and isotopic composition of tektites and impact glasses: Constraints on source regions. *Geochimica et Cosmochimica Acta*, 59(3), 613-624.

674 Chen, H., Savage, P. S., Teng, F. Z., Helz, R. T., & Moynier, F. (2013). Zinc isotope fractionation during magmatic differentiation and the isotopic composition of the bulk Earth. *Earth and Planetary Science Letters*, 369, 34-42.
677 Cloquet, C., Carignan, J., Lehmann, M. F., & Vanhaecke, F. (2008). Variation in the isotopic composition of zinc in the natural environment and the use of zinc isotopes in biogeosciences: a review. *Analytical and Bioanalytical Chemistry*, 390(2), 451-463.

680 Davis, A. M., Hashimoto, A., Clayton, R. N., & Mayeda, T. K. (1990). Isotope mass fractionation during evaporation of Mg2SiO4. *Nature*, 347(6294), 655.

682 Day, J. M. D., Moynier, F., & Shearer, C. K. (2017a). Late-stage magmatic outgassing from a volatile-depleted Moon. *Proceedings of the National Academy of Sciences*, 114(36), 9547-9551.

684 Day, J. M. D., Moynier, F., Meshik, A. P., Pradivtseva, O. V., & Petit, D. R. (2017b). Evaporative fractionation of zinc during the first nuclear detonation. *Science advances*, 3(2), e1602668.

686 Day, J. M. D., van Kooten, E. M., Hofmann, B. A., & Moynier, F. (2020a). Mare basalt meteorites, magnesian-suite rocks and KREEP reveal loss of zinc during and after lunar formation. *Earth and Planetary Science Letters*, 531, 115998.

689 Day, J. M. D., Moynier, F., Sossi, P. A., Wang, K., Meshik, A. P., Pradivtseva, O. V., & Pettit, D. R. (2020b). Moderately volatile element behaviour at high temperature determined from nuclear detonation. *Geochemical Perspectives Letters*, 13, 54-60.

692 Dhaliwal, J. K., Day, J. M. D., & Moynier, F. (2018). Volatile element loss during planetary magma ocean phases. *Icarus*, 300, 249-260.

694 Ding, T. P., Thode, H. G., & Rees, C. E. (1983). Sulphur content and sulphur isotope composition of orange and black glasses in Apollo 17 drive tube 74002/1. *Geochimica et Cosmochimica Acta*, 47(3), 491-496.

697 Dong, S., & Wasylenki, L. E. (2016). Zinc isotope fractionation during adsorption to calcite at high and low ionic strength. *Chemical Geology*, 447, 70-78.

699 Eby, G. N., Charnley, N., Pirrie, D., Hermes, R., Smoliga, J., & Rollinson, G. (2015). Trinitite redux: Mineralogy and petrology. *American Mineralogist*, 100(2-3), 427-441.

701 Elardo, S. M., Draper, D. S., & Shearer Jr, C. K. (2011). Lunar Magma Ocean crystallization revisited: Bulk composition, early cumulate mineralogy, and the source regions of the highlands Mg-suite. *Geochimica et Cosmochimica Acta*, 75(11), 3024-3045.

704 Elkins-Tanton, L. T., Van Orman, J. A., Hager, B. H., & Grove, T. L. (2002). Re-examination of the lunar magma ocean cumulate overturn hypothesis: melting or mixing is required. *Earth and Planetary Science Letters*, 196(3-4), 239-249.

707 Eppich, G. R., Knight, K. B., Jacomb-Hood, T. W., Spriggs, G. D., & Hutcheon, I. D. (2014). Constraints on fallout melt glass formation from a near-surface nuclear test. *Journal of Radioanalytical and Nuclear Chemistry*, 302(1), 593-609.

710 Feng, L. P., Zhou, L., Liu, J., Hu, Z. C., & Liu, Y. S. (2019). Determination of Gallium Isotopic Compositions in Reference Materials. *Geostandards and Geoanalytical Research*.
712 Fernandez, A., & Borrok, D. M. (2009). Fractionation of Cu, Fe, and Zn isotopes during the oxidative weathering of sulfide-rich rocks. *Chemical Geology*, 264(1-4), 1-12.

714 French, B. M. (1998). Traces of catastrophe: A handbook of shock-metamorphic effects in terrestrial meteorite impact structures.

716 Glass, B. P., & Koeberl, C. (2006). Australasian microtektites and associated impact ejecta in the South China Sea and the Middle Pleistocene supereruption of Toba. *Meteoritics & Planetary Science*, 41(2), 305-326.

719 Glasstone, S., & Dolan, P. J. (1977). The effects of nuclear weapons (No. TID-28061). Department of Defense, Washington, DC (USA); Department of Energy, Washington, DC (USA).

721 Grossman, L., Ebel, D. S., Simon, S. B., Davis, A. M., Richter, F. M., & Parsad, N. M. (2000). Major element chemical and isotopic compositions of refractory inclusions in C3 chondrites: The separate roles of condensation and evaporation. *Geochimica et Cosmochimica Acta*, 64(16), 2879-2894.

724 Grossman, L., Ebel, D. S., & Simon, S. B. (2002). Formation of refractory inclusions by evaporation of condensate precursors. *Geochimica et Cosmochimica Acta*, 66(1), 145-161.

726 Guinoiseau, D., Gélabert, A., Moureau, J., Louvat, P., & Benedetti, M. F. (2016). Zn isotope fractionation during sorption onto kaolinite. *Environmental science & technology*, 50(4), 1844-1852.

728 Herzog, G. F., Moynier, F., Albarède, F., & Berezhnoy, A. A. (2009). Isotopic and elemental abundances of copper and zinc in lunar samples, Zagami, Pele's hairs, and a terrestrial basalt. *Geochimica et Cosmochimica Acta*, 73(19), 5884-5904.

731 Humayun M. (2018) Gallium. In: White W.M. (eds) *Encyclopedia of Geochemistry*. Encyclopedia of Earth Sciences Series. Springer, Cham

733 Howard, K. T. (2011). Volatile enhanced dispersal of high velocity impact melts and the origin of tektites. *Proceedings of the Geologists' Association*, 122(3), 363-382.

735 John, S. G., Park, J. G., Zhang, Z., & Boyle, E. A. (2007). The isotopic composition of some common forms of anthropogenic zinc. *Chemical Geology*, 245(1-2), 61-69.

737 Jones, A. P., Genge, M., & Carmody, L. (2013). Carbonate melts and carbonatites. *Rev Mineral Geochem*, 75(1), 289-322.

739 Kato, C., Moynier, F., Valdes, M. C., Dhaliwal, J. K., & Day, J. M. D., (2015). Extensive volatile loss during formation and differentiation of the Moon. *Nature communications*, 6, 7617.

741 Kato, C., & Moynier, F. (2017). Gallium isotopic evidence for extensive volatile loss from the Moon during its formation. *Science advances*, 3(7), e1700571.

743 Kato, C., Moynier, F., Foriel, J., Teng, F. Z., & Puchtel, I. S. (2017). The gallium isotopic composition of the bulk silicate Earth. *Chemical Geology*, 448, 164-172.

745 Knight, K. B., Kita, N. T., Mendybaev, R. A., Richter, F. M., Davis, A. M., & Valley, J. W. (2009). Silicon isotopic fractionation of CAI-like vacuum evaporation residues. *Geochimica et Cosmochimica Acta*, 73(20), 6390-6401.
748 Koeberl, C. (1988). The origin of tektites: A geochemical discussion. Proceedings of the NIPR Symposium.

749 260-291.

750 Koeberl, C. (1992). Geochemistry and origin of Muong Nong-type tektites. Geochimica et Cosmochimica

751 Acta, 56(3), 1033-1064.

752 Le Roux, V., Dasgupta, R., & Lee, C. T. A. (2015). Recommended mineral-melt partition coefficients for

753 FRTEs (Cu), Ga, and Ge during mantle melting. American Mineralogist, 100(11-12), 2533-2544.

754 Lewis, L.A., Knight, K.B., Matzel, J.E., Prussin, S.G., Zimmer, M.M., Kinman, W.S., Ryerson, F.J. and

755 Hutcheon, I.D., 2015. Spatially-resolved analyses of aerodynamic fallout from a uranium-fueled nuclear

756 test. Journal of environmental radioactivity, 148, 183-195.

757 Lock, S. J., Stewart, S. T., Petaev, M. I., Leinhardt, Z., Mace, M. T., Jacobsen, S. B., & Cuk, M. (2018). The

758 origin of the Moon within a terrestrial synestia. Journal of Geophysical Research: Planets, 123(4), 910-

759 951.

760 Lodders, K. (2003). Solar system abundances and condensation temperatures of the elements. The

761 Astrophysical Journal, 591(2), 1220.

762 Lv, Y., Liu, S. A., Zhu, J. M., & Li, S. (2016). Copper and zinc isotope fractionation during deposition and

763 weathering of highly metalliferous black shales in central China. Chemical Geology, 445, 24-35.

764 Ma, P., Aggrey, K., Tonzola, C., Schnabel, C., De Nicola, P., Herzog, G.F., Wasson, J.T., Glass, B.P., Brown,

765 L., Tera, F. and Middleton, R., 2004. Beryllium-10 in Australasian tektites: Constraints on the location of

766 the source crater. Geochimica et Cosmochimica Acta, 68(19), 3883-3896.

767 Macris, C. A., Asimow, P. D., Badro, J., Eiler, J. M., Zhang, Y., & Stolper, E. M. (2018). Seconds after

768 impact: Insights into the thermal history of impact ejecta from diffusion between lechatelierite and host

769 glass in tektites and experiments. Geochimica et Cosmochimica Acta, 241, 69-94.

770 Magna, T., Deutsch, A., Mezger, K., Skála, R., Seitz, H.M., Mizera, J., Řanda, Z. and Adolph, L., 2011.

771 Lithium in tektites and impact glasses: Implications for sources, histories and large impacts. Geochimica

772 et Cosmochimica Acta, 75(8), 2137-2158.

773 Malvin, D. J., & Drake, M. J. (1987). Experimental determination of crystal/melt partitioning of Ga and

774 Ge in the system forsterite-anorthite-diopside. Geochimica et Cosmochimica Acta, 51(8), 2117-2128.

775 Mavromatis, V., González, A. G., Dietzel, M., & Schott, J. (2019). Zinc isotope fractionation during the

776 inorganic precipitation of calcite–Towards a new pH proxy. Geochimica et Cosmochimica Acta, 244, 99-

777 112.

778 McDonough, W. F., & Sun, S. S. (1995). The composition of the Earth. *Chemical geology*, 120(3-4), 223-

779 253.

780 McKenzie, D. (1985). The extraction of magma from the crust and mantle. Earth and Planetary Science

781 Letters, 74(1), 81-91.

782 Mendybaev, R. A., Williams, C. D., Spicuzza, M. J., Richter, F. M., Valley, J. W., Fedkin, A. V., & Wadhwa,

783 M. (2017). Thermal and chemical evolution in the early Solar System as recorded by FUN CAIs: Part II–
784 Laboratory evaporation of potential CMS-1 precursor material. Geochimica et Cosmochimica Acta, 201,
785 49-64.

786 Mizera, J., Řanda, Z., & Kameník, J. (2016). On a possible parent crater for Australasian tektites:

787 Geochemical, isotopic, geographical and other constraints. Earth-science reviews, 154, 123-137.

788 Moynier, F., Albarède, F., & Herzog, G. F. (2006). Isotopic composition of zinc, copper, and iron in lunar

789 samples. *Geochimica et Cosmochimica Acta*, **70**(24), 6103-6117.

790 Moynier, F., Beck, P., Jourdan, F., Yin, Q. Z., Reimold, U., & Koeberl, C. (2009). Isotopic fractionation of

791 zinc in tektites. Earth and Planetary Science Letters, 277(3-4), 482-489.

792 Newsom, H. E., & Taylor, S. R. (1989). Geochemical implications of the formation of the Moon by a single

793 giant impact. Nature, 338(6210), 29-34.

794 Nie, N. X., & Dauphas, N. (2019). Vapor Drainage in the Protolunar Disk as the Cause for the Depletion in

795 Volatile Elements of the Moon. The Astrophysical Journal Letters, 884(2), L48.

796 Norris, C. A., & Wood, B. J. (2017). Earth's volatile contents established by melting and vaporization.

797 Nature, 549(7673), 507-510.

798 Nyquist, L. E., Wooden, J. L., Shih, C. Y., Wiesmann, H., & Bansal, B. M. (1981). Isotopic and REE studies

799 of lunar basalt 12038: implications for petrogenesis of aluminous mare basalts. Earth and Planetary

800 Science Letters, 55(3), 335-355.

801 O'Neill, H. S. C. (1991). The origin of the Moon and the early history of the Earth—A chemical model.

802 Part 1: The Moon. Geochimica et Cosmochimica Acta, 55(4), 1135-1157.

803 Osinski, G. R., & Spray, J. G. (2001). Impact-generated carbonate melts: evidence from the Haughton

804 structure, Canada. Earth and Planetary Science Letters, 194(1-2), 17-29.

805 Paniello, R. C., Day, J. M., & Moynier, F. (2012). Zinc isotopic evidence for the origin of the Moon.

806 Nature, 490(7420), 376-379.

807 Papanastassiou, D. A., & Wasserburg, G. J. (1971). Lunar chronology and evolution from Rb-Sr studies of

808 Apollo 11 and 12 samples. Earth and Planetary Science Letters, 11(1-5), 37-62.

809 Petit, J. C., De Jong, J., Chou, L., & Mattielli, N. (2008). Development of Cu and Zn Isotope MC-ICP-MS

810 Measurements: Application to Suspended Particulate Matter and Sediments from the Scheldt Estuary.

811 Geostandards and Geoanalytical Research, 32(2), 149-166.

812 Pokrovsky, O. S., Viers, J., & Freydier, R. (2005). Zinc stable isotope fractionation during its adsorption on

813 oxides and hydroxides. Journal of Colloid and Interface Science, 291(1), 192-200.

814 Pringle, E. A., & Moynier, F. (2017). Rubidium isotopic composition of the Earth, meteorites, and the

815 Moon: Evidence for the origin of volatile loss during planetary accretion. Earth and Planetary Science

816 Letters, 473, 62-70.

817 Renggli, C. J., King, P. L., Henley, R. W., & Norman, M. D. (2017). Volcanic gas composition, metal

818 dispersion and deposition during explosive volcanic eruptions on the Moon. Geochimica et

819 Cosmochimica Acta, 206, 296-311.
Richter, F. M., Davis, A. M., Ebel, D. S., & Hashimoto, A. (2002). Elemental and isotopic fractionation of Type B calcium-, aluminum-rich inclusions: experiments, theoretical considerations, and constraints on their thermal evolution. *Geochimica et Cosmochimica Acta*, 66(3), 521-540.

Richter, F. M. (2004). Timescales determining the degree of kinetic isotope fractionation by evaporation and condensation. *Geochimica et Cosmochimica acta*, 68(23), 4971-4992.

Richter, F. M., Janney, P. E., Mendybaev, R. A., Davis, A. M., & Wadhwa, M. (2007). Elemental and isotopic fractionation of Type B CAI-like liquids by evaporation. *Geochimica et Cosmochimica Acta*, 71(22), 5544-5564.

Richter, F. M., Mendybaev, R. A., Christensen, J. N., Ebel, D., & Gaffney, A. (2011). Laboratory experiments bearing on the origin and evolution of olivine-rich chondrules. *Meteoritics & Planetary Science*, 46(8), 1152-1178.

Ringwood, A. E., & Kesson, S. E. (1977). Basaltic magmatism and the bulk composition of the moon. The Moon, 16(4), 425-464.

Roberts, S. E., Sheffer, A. A., McCanta, M. C., Dyar, M. D., & Sklute, E. C. (2019). Oxidation state of iron in fulgurites and Trinitite: Implications for redox changes during abrupt high-temperature and pressure events. *Geochimica et Cosmochimica Acta*, 266, 332-350.

Rodovská, Z., Magna, T., Žák, K., Kato, C., Savage, P.S., Moynier, F., Skála, R. and Ježek, J., 2017. Implications for behavior of volatile elements during impacts—Zinc and copper systematics in sediments from the Ries impact structure and central European tektites. *Meteoritics & Planetary Science*, 52(10), 2178-2192.

Rudnick, R. L., & Gao, S. (2003). Composition of the continental crust. *Treatise on geochemistry*, 3, 1-64.

Ryder, G. (1990). Lunar samples, lunar accretion and the early bombardment of the Moon. *Eos*, 81(10), 313-323.

Schnetzler, C. C. (1992). Mechanism of Muong Nong-type tektite formation and speculation on the source of Australasian tektites. *Meteoritics*, 27(2), 154-165.

Schwarz, W.H., Trieloff, M., Bollinger, K., Gantert, N., Fernandes, V.A., Meyer, H.P., Povenmire, H., Jessberger, E.K., Guglielmino, M. and Koeberl, C., 2016. Coeval ages of Australasian, Central American and Western Canadian tektites reveal multiple impacts 790 ka ago. *Geochimica et Cosmochimica Acta*, 178, 307-319.

Shearer, C. K., & Papike, J. J. (1999). Magmatic evolution of the Moon. *American Mineralogist*, 84(10), 1469-1494.

Shearer, C.K., Hess, P.C., Wieczorek, M.A., Pritchard, M.E., Parmentier, E.M., Borg, L.E., Longhi, J., Elkins-Tanton, L.T., Neal, C.R., Antonenko, I. and Canup, R.M., 2006. Thermal and magmatic evolution of the Moon. *Reviews in Mineralogy and Geochemistry*, 60(1), 365-518.

Shearer, C. K., Sharp, Z. D., Burger, P. V., McCubbin, F. M., Provencio, P. P., Brearley, A. J., & Steele, A. (2014). Chlorine distribution and its isotopic composition in "rusty rock" 66095. *Implications for volatile*
element enrichments of "rusty rock" and lunar soils, origin of "rusty" alteration, and volatile element behavior on the Moon. Geochimica et Cosmochimica Acta, 139, 411-433.

Siebert, C., Nägler, T. F., & Kramers, J. D. (2001). Determination of molybdenum isotope fractionation by double-spike multicollector inductively coupled plasma mass spectrometry. *Geochemistry, Geophysics, Geosystems, 2*(7).

Snyder, G. A., Taylor, L. A., & Neal, C. R. (1992). A chemical model for generating the sources of mare basalts: Combined equilibrium and fractional crystallization of the lunar magmasphere. Geochimica et Cosmochimica Acta, 56(10), 3809-3823.

Son, T. H., & Koeberl, C. (2005). Chemical variation within fragments of Australasian tektites. *Meteoritics & Planetary Science*, 40(6), 805-815.

Sossi, P. A., Halverson, G. P., Nebel, O., & Eggins, S. M. (2015). Combined separation of Cu, Fe and Zn from rock matrices and improved analytical protocols for stable isotope determination. *Geostandards and Geoanalytical Research, 39*(2), 129-149

Sossi, P. A., Nebel, O., O'Neill, H. S. C., & Moynier, F. (2018). Zinc isotope composition of the Earth and its behaviour during planetary accretion. *Chemical Geology*, 477, 73-84.

Sossi, P. A., Moynier, F., & Van Zuilen, K. (2018). Volatile loss following cooling and accretion of the Moon revealed by chromium isotopes. *Proceedings of the National Academy of Sciences*, 115(43), 10920-10925.

Stern, S. A. (1999). The lunar atmosphere: History, status, current problems, and context. *Reviews of Geophysics*, 37(4), 453-491.

Taylor, S. R., Sachs, M., & Cherry, R. D. (1961). Studies of tektite composition—I Inverse relationship between SiO₂ and the other major constituents. Geochimica et Cosmochimica Acta, 22(2-4), 155-163.

Taylor, L. A., Patchen, A., Mayne, R. G., & Taylor, D. H. (2004). The most reduced rock from the moon, Apollo 14 basalt 14053: Its unique features and their origin. *American Mineralogist*, 89(11-12), 1617-1624.

Tian, Z., Jolliff, B.L., Korotev, R.L., Fegley Jr, B., Lodders, K., Day, J.M., Chen, H. and Wang, K., 2020.

Potassium isotopic composition of the Moon. Geochimica et Cosmochimica Acta. 280, 263-280.

Touboul, M., Kleine, T., Bourdon, B., Palme, H., & Wieler, R. (2007). Late formation and prolonged differentiation of the Moon inferred from W isotopes in lunar metals. *Nature*, 450(7173), 1206.

Tsuchiyama, A., Tachibana, S., & Takahashi, T. (1999). Evaporation of forsterite in the primordial solar nebula; rates and accompanied isotopic fractionation. Geochimica et cosmochimica acta, 63(16), 2451-2466.

Turner, G. (1970). Argon-40/argon-39 dating of lunar rock samples. *Science*, 167(3918), 466-468.

Wang, J., Davis, A. M., Clayton, R. N., & Hashimoto, A. (1999). Evaporation of single crystal forsterite: Evaporation kinetics, magnesium isotope fractionation, and implications of mass-dependent isotopic fractionation of a diffusion-controlled reservoir. *Geochimica et Cosmochimica Acta*, 63(6), 953-966.
892 Wang, K., & Jacobsen, S. B. (2016). Potassium isotopic evidence for a high-energy giant impact origin of the Moon. *Nature*, **538**(7626), 487-490.

894 Wang, Z.Z., Liu, S.A., Liu, J., Huang, J., Xiao, Y., Chu, Z.Y., Zhao, X.M. and Tang, L., 2017. Zinc isotope fractionation during mantle melting and constraints on the Zn isotope composition of Earth's upper mantle. *Geochimica et Cosmochimica Acta*, **198**, 151-167.

897 Wänke, H., Baddenhausen, H., Dreibus, G., Jagoutz, E., Kruse, H., Palme, H., Spettel, B. and Teschke, F., 1973. Multielement analyses of Apollo 15, 16, and 17 samples and the bulk composition of the moon. In *Lunar and planetary science conference proceedings (Vol. 4, p. 1461)*.

900 Warren, P. H. (1989). KREEP: major-element diversity, trace-element uniformity (almost). In *Moon in transition: Apollo 14, KREEP, and evolved lunar rocks* (pp. 149-153).

902 Weisz, D.G., Jacobsen, B., Marks, N.E., Knight, K.B., Isselhardt, B.H., Matzel, J.E., Weber, P.K., Prussin, S.G. and Hutcheon, I.D., 2017. Deposition of vaporized species onto glassy fallout from a near-surface nuclear test. *Geochimica et cosmochimica acta*, **201**, 410-426.

905 Weisz, D. G., Jacobsen, B., Marks, N. E., Knight, K. B., Isselhardt, B. H., & Matzel, J. E. (2018). Diffusive mass transport in agglomerated glassy fallout from a near-surface nuclear test. *Geochimica et Cosmochimica Acta*, **223**, 377-388.

908 Wetzel, D. T., Hauri, E. H., Saal, A. E., & Rutherford, M. J. (2015). Carbon content and degassing history of the lunar volcanic glasses. *Nature Geoscience*, **8**(10), 755-758.

910 Wiechert, U., Halliday, A. N., Lee, D. C., Snyder, G. A., Taylor, L. A., & Rumble, D. (2001). Oxygen isotopes and the Moon-forming giant impact. *Science*, **294**(5541), 345-348.

912 Wimpenny, J., Marks, N., Knight, K., Rolison, J.M., Borg, L., Eppich, G., Badro, J., Ryerson, F.J., Sanborn, M., Huyskens, M.H. and Yin, Q.Z., 2019. Experimental determination of Zn isotope fractionation during evaporative loss at extreme temperatures. *Geochimica et Cosmochimica Acta*, **259**, 391-411.

915 Young, E.D., Shahar, A., Nimmo, F., Schlichting, H.E., Schauble, E.A., Tang, H. and Labidi, J., 2019. Near-equilibrium isotope fractionation during planetesimal evaporation. *Icarus*, **323**, 1-15.

917 Yuan, W., Chen, J.B., Birck, J.L., Yin, Z.Y., Yuan, S.L., Cai, H.M., Wang, Z.W., Huang, Q. and Wang, Z.H., 2016. Precise analysis of gallium isotopic composition by MC-ICP-MS. *Analytical chemistry*, **88**(19), 9606-9613.

920 Yuan, W., Saldi, G.D., Chen, J., Zuccolini, M.V., Birck, J.L., Liu, Y. and Schott, J., 2018. Gallium isotope fractionation during Ga adsorption on calcite and goethite. *Geochimica et Cosmochimica Acta*, **223**, 350-363.

923 Zhang, J., Dauphas, N., Davis, A. M., Leya, I., & Fedkin, A. (2012). The proto-Earth as a significant source of lunar material. *Nature Geoscience*, **5**(4), 251.

925 Zhang, T., Zhou, L., Yang, L., Wang, Q., Feng, L. P., & Liu, Y. S. (2016). High precision measurements of gallium isotopic compositions in geological materials by MC-ICP-MS. *Journal of Analytical Atomic Spectrometry*, **31**(8), 1673-1679.
928 **Figure captions**

929 **Fig. 1** – Purification scheme for separating Ga from matrix elements. Chemical purification was
930 adapted from Kato et al., (2017) and Yuan et al., (2017).

931 **Fig. 2** – Ga isotope ratio of a) the in-house Ga ICP-MS standard, referred to as LLNL-Ga, and b) USGS rock standards BCR-2 and BIR-1a. All samples are bracketed to the NIST isotope standard SRM 994. Internal uncertainties are typically between 0.01-0.03‰ (2σ).
932
933

934 **Fig. 3** – Standard addition plot, showing the measured δ<sup>71</sup>Ga value of 5 mixtures of BCR-2 and SRM 994 mixed to different proportions and processed through ion exchange chromatography.
935

936 **Fig. 4** – Ga systematics of laser-heated rhyolitic soil, fallout melt glass and Australasian tektites. Typical uncertainties for the δ<sup>71</sup>Ga values are <0.06‰ (2σ).
937

938 **Fig. 5** – Rayleigh fractionation plots for the three sample suites showing the extent of Ga and Zn isotope fractionation as a function of element depletion. Rayleigh fractionation lines of 0.9988 and 0.9997 are plotted for reference. Zn data is a combination of new data and data taken from Wimpenny et al., (2019). All uncertainties on the isotope ratios are 2σ.
939
940
941

942 **Fig. 6** – Calculation of the Ga fractionation factor during evaporative loss from the laser heating experiments. The data are consistent with an α of 0.99891 ± 0.00024. The regression uses the error weighted least squares algorithm of York et al., (2004) and is calculated using IsoplotR (Vermeesh 2018). The outlying data points rhyolite-8 and -10 have been excluded from this regression.
943
944
945
946

947 **Fig. 7** – δ<sup>71</sup>Ga- δ<sup>66</sup>Zn systematics of the laser heating experiments, fallout melt glass and Australasian tektites. The values Δ<sup>71</sup>Ga<sub>R<sub>0</sub>-R</sub> and Δ<sup>66</sup>Zn<sub>R<sub>0</sub>-R</sub> represent the isotopic difference between the volatilized sample (R) and starting material (R<sub>0</sub>). For reference we have plotted lines of slope 1, 0.65 and 0.1. Here, the slope 1 line represents the expected isotopic composition of a volatile depleted sample that has lost equal proportions of Ga and Zn; in this case the δ<sup>71</sup>Ga- δ<sup>66</sup>Zn systematics are simply controlled by the relative mass difference between the two isotope systems (slope = 0.93). The outlying data points rhyolite-8 and -10 plot to the left of the slope 1 line. The lines of slope 0.65 and 0.1 represent Ga loss of 65% and 10% of Zn loss respectively. All uncertainties are 2σ.
948
949
950
951
952
953
954
955

956 **Fig. 8** – δ<sup>71</sup>Ga- δ<sup>66</sup>Zn systematics of mare basalts (Kato et al., 2015; Kato & Moynier 2017). The values δ<sup>71</sup>Ga<sub>BSE</sub> and δ<sup>66</sup>Zn<sub>BSE</sub> represent the isotopic difference between the mare basalts and the BSE. For reference we have plotted lines of slope 1, 0.65 and 0.1, where the slope 1 line represents the expected isotopic composition of a volatile depleted sample that has lost equal proportions of Ga and Zn. The lines of slope 0.65 and 0.1 represent Ga losses of 65% and 10% of
961 total Zn loss respectively. All uncertainties are 2σ. All lunar data is taken from Kato et al.,

962 (2015) and Kato & Moyier (2017). For clarity, the isotopic data for 14053 and 12005, which have

963 negative δ<sup>66</sup>Zn values, are not shown.

964

965
<table><caption>**Table 1** – Gallium purification scheme by ion-exchange chromatography.</caption><thead><tr><th></th><th>Reagent</th><th>Volume</th><th>Eluted element</th></tr></thead><tbody><tr><td colspan="4"><strong>Column 1 - Separate gallium from sample matrix</strong></td></tr><tr><td>Load 1.5 ml of Bio-Rad AG1-X8 (200-400 mesh) anion resin into Bio-Rad 2ml Polyprep columns</td><td></td><td></td><td></td></tr><tr><td>Clean resin</td><td>MQ H2O</td><td>10 ml</td><td></td></tr><tr><td></td><td>0.5 M HCl</td><td>10 ml</td><td></td></tr><tr><td></td><td>6 M HCl</td><td>6 ml</td><td></td></tr><tr><td>Pre-condition resin</td><td>6 M HCl</td><td>0.5 ml</td><td></td></tr><tr><td>Dissolve and load sample. Place beaker to collect sample matrix</td><td>6 M HCl</td><td>10 ml</td><td>Mg, Al, Na, REE's, most matrix</td></tr><tr><td>Elute matrix</td><td>6 M HCl</td><td>10 ml</td><td></td></tr><tr><td>Collect gallium</td><td>0.5 M HCl</td><td>8 ml</td><td>Ga, Fe, U, Mo</td></tr><tr><td>Column and resin are discarded after use</td><td></td><td></td><td></td></tr><tr><td colspan="4"><strong>Column 2 - Separate gallium from iron</strong></td></tr><tr><td>Load 2 ml of Eichrom LN-spec (100-150 µm) resin into Bio-Rad 2ml Polyprep columns</td><td></td><td></td><td></td></tr><tr><td>Clean resin</td><td>MQ H2O</td><td>10 ml</td><td></td></tr><tr><td></td><td>6 M HCl</td><td>10 ml</td><td></td></tr><tr><td></td><td>1 M HCl</td><td>6 ml</td><td></td></tr><tr><td>Pre-condition</td><td>1 M HCl</td><td>6.5 ml</td><td>Ga</td></tr><tr><td>Load and collect gallium</td><td>1 M HCl</td><td>6.5 ml</td><td></td></tr><tr><td>Column and resin are discarded after use</td><td></td><td></td><td></td></tr><tr><td colspan="4"><strong>Column 3 - Clean up gallium from any residual iron and trace matrix elements</strong></td></tr><tr><td>1ml of Bio-Rad AG50-X12 (200-400 mesh) resin loaded into Bio-Rad 2ml glass Econo columns</td><td></td><td></td><td></td></tr><tr><td>Clean</td><td>MQ H2O</td><td>10 ml</td><td></td></tr><tr><td></td><td>6 M HCl</td><td>10 ml</td><td></td></tr><tr><td></td><td>2 M HCl</td><td>6 ml</td><td></td></tr><tr><td>Pre-condition</td><td>2 M HCl</td><td>0.5 ml</td><td></td></tr><tr><td>Load sample</td><td>2 M HCl</td><td>0.5 ml</td><td></td></tr><tr><td>Elute Matrix</td><td>2 M HCl</td><td>7.5 ml</td><td>Fe, matrix including Cu, Mo, Ba</td></tr><tr><td>Collect Ga</td><td>2 M HCl</td><td>12 ml</td><td>Ga</td></tr><tr><td>Clean resin</td><td>6 M HCl</td><td>10 ml</td><td></td></tr><tr><td></td><td>MQ H2O</td><td>10 ml</td><td></td></tr><tr><td>Store column for future use</td><td></td><td></td><td></td></tr></tbody></table>
<table><caption>**Table 2** – Results of gallium and zinc isotope analyses of laser heated rhyolitic soil, nuclear fallout glass and australite tektites. Isotopic compositions of the BSE for Ga and Zn are taken from Kato et al., (2017) and Sossi et al., (2018). Zn isotope data that was previously published in Wimpenny et al., (2019) is italicized.</caption><thead><tr><th></th><th>Heating time (s)</th><th>Max T (°C)</th><th>δ<sup>1</sup>Ga (SRM 994)</th><th>δ<sup>1</sup>Ga (BSE)</th><th>Int Uncert.</th><th>δ<sup>86</sup>Zn (IRMM-3702)</th><th>δ<sup>86</sup>Zn (JMC-Lyon)</th><th>Int Uncert.</th><th>Ga (ppm)</th><th>2σ</th><th>Zn (ppm)</th><th>2σ</th><th>Al (ppm)</th><th>2σ</th></tr></thead><tbody><tr><td>Rhyolite soil-bulk</td><td>N/A</td><td>N/A</td><td>0.47</td><td>-0.03</td><td>0.07</td><td>-0.02</td><td>0.28</td><td>0.02</td><td>29.07</td><td>2.97</td><td>153.0</td><td>6.0</td><td>67200</td><td>2300</td></tr><tr><td>Rhyolite soil-17</td><td>120</td><td>1600</td><td>0.49</td><td>-0.01</td><td>0.07</td><td>0.10</td><td>0.40</td><td>0.02</td><td>30.87</td><td>0.79</td><td>149.5</td><td>5.3</td><td>69100</td><td>900</td></tr><tr><td>Rhyolite soil-3</td><td>10</td><td>1800</td><td>0.57</td><td>0.07</td><td>0.02</td><td></td><td></td><td></td><td>30.76</td><td>0.96</td><td>121.2</td><td>4.5</td><td>68700</td><td>2100</td></tr><tr><td>Rhyolite soil-4</td><td>120</td><td>1800</td><td>0.61</td><td>0.11</td><td>0.05</td><td>0.24</td><td>0.54</td><td>0.02</td><td>29.04</td><td>0.39</td><td>122.5</td><td>1.9</td><td>69500</td><td>1800</td></tr><tr><td>Rhyolite soil-11</td><td>5</td><td>2000</td><td>0.62</td><td>0.12</td><td>0.02</td><td>0.21</td><td>0.51</td><td>0.02</td><td>30.66</td><td>0.70</td><td>135.6</td><td>3.0</td><td>69600</td><td>1500</td></tr><tr><td>Rhyolite soil-8</td><td>120</td><td>2000</td><td>1.18</td><td>0.68</td><td>0.04</td><td>0.55</td><td>0.85</td><td>0.02</td><td>25.68</td><td>0.81</td><td>102.7</td><td>2.6</td><td>72200</td><td>2000</td></tr><tr><td>Rhyolite soil-10</td><td>5</td><td>2200</td><td>1.12</td><td>0.62</td><td>0.07</td><td>0.40</td><td>0.70</td><td>0.02</td><td>28.18</td><td>0.37</td><td>116.5</td><td>1.2</td><td>69100</td><td>1500</td></tr><tr><td>Rhyolite soil-6</td><td>10</td><td>2200</td><td>0.76</td><td>0.26</td><td>0.07</td><td>0.49</td><td>0.79</td><td>0.03</td><td>26.92</td><td>0.83</td><td>117.7</td><td>3.7</td><td>73000</td><td>1200</td></tr><tr><td>Rhyolite soil-15</td><td>60</td><td>2200</td><td>1.30</td><td>0.80</td><td>0.07</td><td>1.19</td><td>1.49</td><td>0.02</td><td>19.26</td><td>0.30</td><td>83.3</td><td>3.0</td><td>85700</td><td>1200</td></tr><tr><td>Rhyolite soil-15 rpt</td><td></td><td></td><td>1.25</td><td>0.75</td><td>0.02</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Fallout-1</td><td></td><td></td><td></td><td></td><td></td><td>0.29</td><td>0.59</td><td>0.02</td><td></td><td></td><td>45.77</td><td>1.1</td><td>63000</td><td>800</td></tr><tr><td>Fallout-2</td><td></td><td></td><td></td><td></td><td></td><td>0.03</td><td>0.33</td><td>0.02</td><td></td><td></td><td>77.22</td><td>2.6</td><td>67600</td><td>700</td></tr><tr><td>Fallout-3</td><td></td><td></td><td>0.85</td><td>0.35</td><td>0.06</td><td>0.17</td><td>0.47</td><td>0.02</td><td>12.84</td><td>0.27</td><td>86.89</td><td>3.3</td><td>63000</td><td>800</td></tr><tr><td>Fallout-4</td><td></td><td></td><td>0.49</td><td>-0.01</td><td>0.02</td><td>0.37</td><td>0.67</td><td>0.02</td><td>21.07</td><td>0.67</td><td>55.81</td><td>1.1</td><td>67600</td><td>700</td></tr><tr><td>Fallout-5</td><td></td><td></td><td>0.67</td><td>0.17</td><td>0.04</td><td>0.10</td><td>0.40</td><td>0.02</td><td>16.28</td><td>0.62</td><td>89.99</td><td>1.6</td><td>60400</td><td>1300</td></tr><tr><td>Fallout-6</td><td></td><td></td><td>0.79</td><td>0.29</td><td>0.05</td><td>0.03</td><td>0.33</td><td>0.02</td><td>15.23</td><td>0.28</td><td>79.22</td><td>2.3</td><td>59100</td><td>1300</td></tr><tr><td>Fallout-7</td><td></td><td></td><td>0.49</td><td>-0.01</td><td>0.08</td><td>0.10</td><td>0.40</td><td>0.02</td><td>24.17</td><td>0.46</td><td>75.09</td><td>1.9</td><td>59600</td><td>1600</td></tr><tr><td>Australite-2019-2</td><td></td><td></td><td>0.59</td><td>0.09</td><td>0.02</td><td>1.31</td><td>1.61</td><td>0.02</td><td>8.39</td><td>0.16</td><td>11.95</td><td>0.18</td><td>62200</td><td>1200</td></tr><tr><td>Australite-2019-3</td><td></td><td></td><td>0.61</td><td>0.11</td><td>0.02</td><td>1.18</td><td>1.48</td><td>0.02</td><td>8.44</td><td>0.24</td><td>13.53</td><td>0.49</td><td>61600</td><td>1000</td></tr><tr><td>Australite-2019-4</td><td></td><td></td><td>0.83</td><td>0.33</td><td>0.02</td><td>1.10</td><td>1.40</td><td>0.01</td><td>6.16</td><td>0.09</td><td>12.98</td><td>0.31</td><td>56000</td><td>1300</td></tr><tr><td>Australite-2019-5</td><td></td><td></td><td>0.88</td><td>0.38</td><td>0.02</td><td>1.06</td><td>1.36</td><td>0.01</td><td>5.95</td><td>0.03</td><td>13.01</td><td>0.63</td><td>55200</td><td>2700</td></tr><tr><td>Australite-2019-6</td><td></td><td></td><td>0.67</td><td>0.17</td><td>0.01</td><td>1.83</td><td>2.13</td><td>0.01</td><td>6.96</td><td>0.10</td><td>7.59</td><td>0.21</td><td>64400</td><td>1300</td></tr><tr><td>Australite-3</td><td></td><td></td><td>0.87</td><td>0.37</td><td>0.02</td><td>0.89</td><td>1.19</td><td>0.01</td><td>5.98</td><td>0.12</td><td>20.00</td><td>1.28</td><td>53900</td><td>200</td></tr><tr><td>Australite-4</td><td></td><td></td><td>0.67</td><td>0.17</td><td>0.02</td><td>1.73</td><td>2.03</td><td>0.02</td><td>6.99</td><td>0.17</td><td>14.65</td><td>0.37</td><td>61700</td><td>800</td></tr><tr><td>Australite-5</td><td></td><td></td><td>0.73</td><td>0.23</td><td>0.02</td><td></td><td></td><td></td><td>7.26</td><td>0.17</td><td>11.96</td><td>0.85</td><td>67000</td><td>800</td></tr><tr><td>Australite-9</td><td></td><td></td><td>0.54</td><td>0.04</td><td>0.03</td><td></td><td></td><td></td><td>8.66</td><td>0.53</td><td>28.19</td><td>1.94</td><td>67500</td><td>600</td></tr><tr><td>BCR-2 (n = 18)</td><td></td><td></td><td>0.52</td><td>0.02</td><td>0.07</td><td></td><td></td><td></td><td>23.1</td><td>0.8</td><td>135.4</td><td>4.7</td><td>66900</td><td>1800</td></tr><tr><td>BIR-1a (n = 7)</td><td></td><td></td><td>0.49</td><td>-0.01</td><td>0.04</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>BHVO-2</td><td></td><td></td><td>0.48</td><td>-0.02</td><td>0.02</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>AGV-2</td><td></td><td></td><td>0.52</td><td>0.02</td><td>0.02</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>RGM-2</td><td></td><td></td><td>0.53</td><td>0.03</td><td>0.02</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>GSP-1</td><td></td><td></td><td>0.53</td><td>0.03</td><td>0.02</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>SDC-1</td><td></td><td></td><td>0.54</td><td>0.04</td><td>0.02</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Syenite-1</td><td></td><td></td><td>0.54</td><td>0.04</td><td>0.02</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>BSE</td><td></td><td></td><td>0.50</td><td>0.00</td><td>0.06</td><td>0.00</td><td>0.16</td><td>0.06</td><td>17.5</td><td>0.7</td><td>67</td><td>6</td><td>81500</td><td>4000</td></tr></tbody></table>

968

969
<figure><img src="image_3.png" /><figcaption>Column 2: 2ml of Eichrom LN-Spec</figcaption></figure>

<figure><img src="image_4.png" /><figcaption>Column 3: 1ml of Bio-Rad AG50-X12 (200-400 mesh)</figcaption></figure>

<figure><img src="image_5.png" /></figure>
<figure><img src="image_6.png" /><figcaption>Fig. 2</figcaption></figure>
976

<figure><img src="image_7.png" /><figcaption>Fig. 3</figcaption></figure>

977

978

979

<figure><img src="image_8.png" /><figcaption>Fig. 4</figcaption></figure>

980

981
a) Rhyolitic laser spheres

<figure><img src="image_9.png" /></figure>

b) Nuclear fallout glass

<figure><img src="image_10.png" /></figure>

c) Australite tektites

<figure><img src="image_11.png" /></figure>
985

<figure><img src="image_12.png" /><figcaption>Fig. 6</figcaption></figure>

986

987
<figure><img src="image_13.png" /><figcaption>Fig. 7</figcaption></figure>
991

<figure><img src="image_14.png" /><figcaption>Fig. 8</figcaption></figure>

992

993