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# Partitioning and isotopic fractionation of Li between mineral phases and alkaline to calc-alkaline melts of explosive and effusive eruptions

Julia Neukampf, Ben S Ellis, Tomas Magna, Oscar Laurent, Yves Marrocchi

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## To cite this version:

Julia Neukampf, Ben S Ellis, Tomas Magna, Oscar Laurent, Yves Marrocchi. Partitioning and isotopic fractionation of Li between mineral phases and alkaline to calc-alkaline melts of explosive and effusive eruptions. Chemical Geology, 2023, 636, pp.121628. 10.1016/j.chemgeo.2023.121628. hal-04931886

HAL Id: hal-04931886

https://ut3-toulouseinp.hal.science/hal-04931886v1

Submitted on 6 Feb 2025

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# Partitioning and isotopic fractionation of Li between mineral phases and alkaline to calc-alkaline melts of explosive and effusive eruptions

Julia Neukampf<sup>a, b</sup>, Ben S. Ellis<sup>a, *</sup>, Tomas Magna<sup>c</sup>, Oscar Laurent<sup>d</sup>, Yves Marrocchi<sup>b</sup>

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## ARTICLE INFO

Editor: Claudia Romano

**Keywords:**

Lithium contents

Partition coefficient

Lithium isotopes

Isotope fractionation, Melt inclusions

## ABSTRACT

Evaluating the lithium (Li) budget of alkaline to calc-alkaline melts and their co-existing mineral phases is crucial to understanding which magmatic or post-eruptive processes can influence the Li contents and isotopic compositions of volcanic deposits. This is also important to be able to better constrain how economically relevant Li deposits form. The lithium contents of many different eruptions from different tectonic settings (intraplate, subduction zone, extensional environment) and cooling environments (quenched, quickly cooled and slowly cooled) were studied by combining new and literature data that were obtained using LA-ICP-MS. We found that the Li concentrations in melt inclusions and re-entrants are independent of tectonic setting (hotspot, subduction zone and rifting) and rarely exceed 100 ppm. When melt inclusions are re-homogenised the Li concentration increases which can lead to the overestimation of the Li budget of a magma.

The highest Li concentrations in the alkaline suite of samples were found in groundmass glass (reaching up to 318 ppm). The Li contents in the co-existing mineral phases decrease in the following order: biotite (0.8–99.3 ppm), plagioclase (0.9–25.1 ppm), clinopyroxene (0.6–31.0 ppm) and sanidine (0.5–38.7 ppm). The highest Li concentrations in the calc-alkaline suite of samples were found in the biotite crystals (10.2–2314 ppm) followed by the groundmass glass (2.1–176 ppm). The Li contents are lower in the other mineral phases, from plagioclase (1.4–35.0 ppm), quartz (3.2–21.1 ppm), olivine (1.0–16.3 ppm), sanidine (0.3–8.1 ppm) to clinopyroxene (0.7–21.8 ppm). The Li concentrations in groundmass glass and co-existing crystal phases are independent of the tectonic setting but appear to be more closely related to their cooling environment. To evaluate the influence of the different cooling environments the Li isotopic compositions (per mil deviation of 7Li/6Li ratio relative to L-SVEC reference material) of different eruptions were also determined. The Li isotopic compositions of groundmass glass and mineral phases reveal a range in δ7Li composition spanning up to 6 ‰ for alkaline samples and >10‰ between bulk sample and mineral phases for calc-alkaline samples. The cooling environment of the samples has a large impact on the Li isotopic compositions of the mineral phases with only quartz seemingly unaffected. Rapidly quenched samples may retain their magmatic δ7Li compositions or will only experience small amounts of diffusional overprinting (e.g., by degassing) whereas slowly cooled samples will experience a prolonged diffusional exchange between minerals and matrix overprinting their magmatic δ7Li compositions.

The Li inventory in quartz crystals may be influenced by the tectonic setting and water concentration in the crystals. To further assess the influence of tectonic settings on Li behaviour we used a large data set which is based on the 13 alkaline to calc-alkaline eruptions from different tectonic settings and calculated the apparent partition coefficients (Kds) between groundmass glass and co-existing mineral phases. Biotite crystals (alkaline systems) return the highest Kds (subduction-extensional setting: 0.01–0.57, oceanic hotspot: 0.17–2.65) whereas sanidine crystals yield the lowest Kds. Sanidines from the subduction-extensional setting (alkaline trend) have the lowest Kds (0.01–0.07; one outlier ranging from 0.02 to 0.24), followed by the oceanic hotspot setting (alkaline trend, 0.02–0.09) and the extensional setting (calc-alkaline trend; 0.01–0.15). Plagioclase crystals from the subduction-extensional setting (alkaline trend) yield Kds between 0.01 and 0.29, crystals from subduction zone settings (0.06–0.81) and extensional settings (0.09–0.50) have larger ranges in their Kds. The clinopyroxene crystals from the subduction zone settings (calc-alkaline trend) have the largest range in Kds (0.03–0.99),

<footnote><sup>a</sup> Institute of Geochemistry and Petrology, ETH Zurich, NW, Clausiusstrasse 25, CH-8092 Zurich, Switzerland</footnote>

<footnote><sup>b</sup> CNRS, Centre de Recherches Pétrographiques et Géochimiques, 15 rue Notre Dame des Pauvres BP 20, F-54500 Vandœuvre, les Nancy, France</footnote>

<footnote><sup>c</sup> Czech Geological Survey, Klářov 3, CZ-11821 Prague 1, Czech Republic</footnote>

<footnote><sup>d</sup> CNRS, Géosciences Environnement Toulouse, Observatoire Midi-Pyrénées, 14 avenue Édouard Belin, F-31400 Toulouse, France</footnote>

<footnote>* Corresponding author.</footnote>

<footnote>E-mail address: julia.neukampf@univ-lorraine.fr (J. Neukampf).</footnote>
followed by the oceanic hotspot setting (alkaline trend, 0.06–0.48) and subduction-extensional setting (alkaline trend, 0.01–0.22). Quartz crystals (only from the calc-alkaline trend) from subduction zone settings have lower Kds (0.09–0.13) compared to the extensional settings (Bishop Tuff: 0.14–0.28; Caetano Tuff: 0.73–0.94). The wide variety of processes that can affect final Li abundances in melt and crystals (degassing, scavenging into fluids, intra-crystal diffusion, hydration of glasses) makes estimating the partition coefficients particularly difficult and results in the wide ranges observed here. Targeted experimental work would thus help to better constrain the behaviour of Li in alkaline compositions.

# 1. Introduction

The lithium (Li) abundances and isotope compositions in igneous systems have received increasing attention during the last two decades as a geochemical tracer of different reservoirs and processes (e.g., Tomascak et al., 2016) such as crustal recycling (e.g., Teng et al., 2004; Magna et al., 2010), cooling (Coogan et al., 2005; Beck et al., 2006; Ellis et al., 2018), explosive volcanic eruptions and related degassing (Berlo et al., 2004; Kent et al., 2007; Neukampf et al., 2021, 2022), and weathering (e.g., Pistiner and Henderson, 2003; Rudnick et al., 2004; Pogge von Strandmann et al., 2013). Lithium is considered a powerful tracer of these processes because it is a moderately incompatible fluid mobile trace element with ~17% relative mass difference between the two stable isotopes ($^{6}$Li and $^{7}$Li) and shows a high diffusivity in melt and minerals (e.g., Richter et al., 2003). The Earth's mantle contains on average ~1.5 ppm Li with $^{67}$Li of ~ +4‰ (e.g., McDonough and Sun, 1995; Seitz et al., 2004; Magna et al., 2006a; Marschall et al., 2017). However, the Earth's mantle is heterogeneous and some mantle-derived peridotites show large variations in their $^{67}$Li values, associated with kinetic effects and/or protracted times of cooling (e.g., Nishio et al., 2004; Jeffcoate et al., 2007; Ionov and Seitz, 2008). Unaltered MORB rocks have ~5 ppm with $^{67}$Li between +1.5‰ and +5.6‰ (e.g., Chan et al., 1992; Elliott et al., 2006; Tomascak et al., 2008; Marschall et al., 2017); however, variations can be found between different MORB segments as a result of kinetic fractionation (Marschall et al., 2017). Due to assimilation, fractional crystallisation and other processes, the Li isotope composition of more evolved melts appears to display significant variations to both lower and higher $^{67}$Li compared with bulk $^{67}$Li mantle and pristine mantle-derived melts (e.g., Ackerman et al., 2015; Hamelin et al., 2009). Although it remains unclear to which extent the individual processes (e.g., crystal fractionation, differentiation, assimilation) contribute to the $^{67}$Li variability in chemically evolved melts (e.g., Teng et al., 2004, 2006; Magna et al., 2010).

Most of the current understanding regarding Li behaviour during igneous processes is based on undifferentiated magmas (e.g., Jeffcoate et al., 2007; Lynn et al., 2018), whilst relatively fewer studies focused on more evolved systems (e.g., Vlastelic et al., 2011; Giuffrida et al., 2018; Ellis et al., 2022c) and even fewer on alkaline compositions (Calderoni et al., 1984; Ackerman et al., 2015). This limited knowledge of the partitioning behaviour of Li between alkaline melts and commonly co-occurring minerals (e.g., Beard et al., 2019; Li et al., 2003) severely hinders the recognition of the fate of Li (and some other metals) in evolved systems. Studying the partitioning behaviour of elements between a crystalline and melt phase (or co-existing fluids) is important to separate the effects of melting, crystallisation from compositionally different melts, and changes in pressure, temperature and geochemical environment. Previous studies identified that during partial melting or crystal fractionation processes Li will preferentially partition into the melt (e.g., Seitz and Woodland, 2000; Ottolini et al., 2009) and that partition coefficients between melts and coexisting mantle mineral phases mostly range between 0.1 and 1.0 (Ryan and Langmuir, 1987; Brenan et al., 1998a, 1998b; Seitz and Woodland, 2000; Seitz et al., 2004; Tomascak et al., 2016; Neukampf et al., 2019). For silica-rich and alkaline compositions, however, partitioning of Li remains comparatively under-explored (Webster et al., 1989; Iveson et al., 2018; Neukampf et al., 2019).

In this study, we focus on the determination of Li abundances and partitioning between melt and mineral phases from a wide compositional range of felsic magmas spanning from alkaline to calc-alkaline in bulk chemistry. We combine new and published data on major and trace elements with a detailed micro-analytical study of phenocrysts and their coexisting groundmass glass from well-characterised localities, coupled with Li isotope measurements of mineral separates. This comprehensive approach allows (i) constraining the major hosts of Li and the condition of formation of economic Li deposits, (ii) understanding how magmatic processes (e.g., degassing, rapid cooling, slow cooling) can influence the isotopic composition, and (iii) deciphering how magmatic processes can be traced using Li.

# 2. Geological backgrounds

The new and literature data presented in this study come from samples that are from globally dispersed and tectonically diverse settings (Fig. 1 and Table 1) and are briefly introduced below. For clarity, we display data as a function of tectonic setting, all data used are provided in supplementary materials on a unit-by-unit basis. The selected deposits and eruption units are from: 1) the Kos-Nisyros Volcanic centre (Greece); 2) the Bishop Tuff (USA); 3) Campi Flegrei (Italy); 4) the Carpenter Ridge Tuff (USA); 5) the Peach Spring Tuff (USA); 6) the McDermitt Tuff (USA); 7) the Green Tuff (Italy); 8) the Ignimbrite 'A' (Spain); 9) the Granadilla Tuff, Poris Tuff and Abrigo Tuff (Spain); 10) the Cuicuiltic Tuff and Xoxoctic Tuff (Mexico); 11) the Caetano tuff

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**Fig. 1.** Overview map of sample locations. Eruptions/volcanoes from subduction zone settings are: 1) Kos-Nisyros volcanic centre, Greece; 2) Carpenter Ridge Tuff, USA; 3) Cuicuiltic Tuff and Xoxoctic Tuff of Los Humeros caldera, Mexico. Eruptions/volcanoes from extensional tectonic setting are 1) Bishop Tuff, USA; 2) Green Tuff from Pantelleria, Italy; 3) Caetano Tuff, USA; 4) Jarbridge Rhyolite, USA. Eruptions/volcanoes from hotspot settings are: 1) Ignimbrite 'A', Gran Canaria; 2) Tuffs Granadilla, Poris and Abrigo from Las Canadas caldera, Tenerife, Canary Islands; 3) McDermitt Tuff, USA; 4) Sete Cidades volcano, São Miguel, Portugal. Eruptions/volcanoes from subduction-extensional settings are: 1) tuffs of Campi Flegrei, Southern Italy and 2) Peach Spring Tuff, USA. The red star represents new data measured for this study. Yellow star indicates locations where literature data was used and yellow star with red fill are eruptions where new data and literature data were combined. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)
(USA); 12) the Pepom tephras (Portugal; 13) the Jarbridge Rhyolite (USA).

## 2.1. Subduction settings

### 2.1.1. Kos-Nisyros Volcanic centre - Greece

The Kos-Nisyros Volcanic centre (KNVC) has been active for the past approximately 3 million years. The rhyolitic Kos Plateau Tuff (KPT; 161 ± 1 ka) is one of the largest eruptions in Europe during the Quaternary.

The KPT is multiply saturated containing sanidine, quartz, plagioclase, biotite, ilmenite, magnetite, zircon, and apatite (Keller, 1969) and represents the eruption of the coolest, wettest magma at the culmination of a caldera cycle (Bachmann et al., 2019). Following the KPT event, the system reverted to more mafic compositions while building the Nisyros stratovolcano. Evolved deposits from Nisyros are dominated by a mineral assemblage variably containing plagioclase, clinopyroxene, orthopyroxene, amphibole, Fe-Ti oxides, and accessory zircon and apatite (Popa et al., 2019). The oldest rhyolitic activity on Nisyros is the Avlaki

<caption>**Table 1**</caption>

Details of the units studied in this paper. Mineralogy key: pl = plagioclase, san = sanidine, anorth = anorthoclase, Afsp = alkali feldspar, qtz = quartz, haü = haüyne, cpx = clinopyroxene, amph = amphibole, bt = biotite, tit = titanite, chev = chevkinite, all = allanite, mt = magnetite, ilm = ilmenite, ap = apatite.

<table><thead><tr><th>Unit name</th><th>Tectonic setting</th><th>Origin</th><th>Age</th><th>Volume</th><th>Composition</th><th>Mineralogy</th><th>Data taken from</th></tr></thead><tbody><tr><td>Kos Plateau Tuff</td><td>subduction</td><td>Kos-Nisyros Volcanic Centre (Greece)</td><td>161 ka</td><td>≥60 km³</td><td>rhyolitic</td><td>san + qtz + pl + bt + ilm + mt + zr + ap<br>pl + cpx + opx + amph + ilm + mt + zr + ap</td><td>Ellis et al. (2022a); this study</td></tr><tr><td>Avlaki lava</td><td>subduction</td><td>Kos-Nisyros Volcanic Centre (Greece)</td><td>118.7 ka</td><td></td><td>rhyolitic</td><td>pl + cpx + opx + amph + ilm + mt + zr + ap</td><td>This study</td></tr><tr><td>Nikia lava</td><td>subduction</td><td>Kos-Nisyros Volcanic Centre (Greece)<br>Long Valley magma chamber (California, USA)</td><td>73.3 ka</td><td>~650 km³</td><td>rhyodacite</td><td>pl + cpx + opx + amph + ilm + mt + zr + ap</td><td>This study</td></tr><tr><td>Bishop Tuff</td><td>extensional</td><td></td><td>0.76 Ma</td><td>km³</td><td>rhyolitic</td><td>san + qtz + pl + mt + bt + ilm + zr + ap</td><td>Ellis et al. (2022a); this study</td></tr><tr><td>Campanian Ignimbrite</td><td>subduction- extensional</td><td>Campi Flegrei caldera (Italy)</td><td>39 ka</td><td>200 km³ (DRE)</td><td>trachyte to phonolite</td><td>san + pl + cpx + bt + mt + ap</td><td>Forni et al. (2016, 2018); Ellis et al. (2022a); this study</td></tr><tr><td>Neapolitan Yellow Tuff</td><td>subduction- extensional</td><td>Campi Flegrei caldera (Italy)</td><td>15 ka</td><td>40 km³ (DRE)</td><td>trachy-andesite to phonolite</td><td>san + pl + cpx + bt + mt + ap</td><td>Forni et al. (2016, 2018); Ellis et al. (2022a); this study</td></tr><tr><td>Astroni</td><td>subduction- extensional</td><td>Campi Flegrei caldera (Italy)</td><td>4.8–3.8 ka</td><td>0.5 km³</td><td>tephriphonolite, phonolite to trachyte</td><td>qtz + san + pl + cpx + opx + bt + amph<br>pl + san + bt + mt + ilm ± cpx ± amph ± zr</td><td>Forni et al. (2016, 2018); Ellis et al. (2022a); this study</td></tr><tr><td>Carpenter Ridge Tuff</td><td>subduction</td><td>Bachelor caldera (Colorado, USA)</td><td>27.55 Ma</td><td>1000 km³</td><td>trachydacite to rhyolite</td><td>san + pl + zr + tit + bt + mt ± ilm ± amph ± cpx ± chev ± all ± qtz</td><td>Lipman and McIntosh (2008); Bachmann et al. (2014); Wolff et al. (2020)</td></tr><tr><td>Peach Spring Tuff</td><td>subduction-extension</td><td>Silver Creek caldera (Arizona-California- Nevada, USA)</td><td>18.8 Ma</td><td>640 km³</td><td>trachyte to rhyolite</td><td>anorth + cpx + fay + mt + ilm ± amph</td><td>Buesch (1992); Ferguson et al. (2013); Pamukcu et al. (2013); Foley et al. (2020)</td></tr><tr><td>McDermitt Tuff</td><td>continental hotspot</td><td>McDermitt caldera (Nevada-Oregon, USA)</td><td>16.39 Ma</td><td>1000 km³</td><td>trachyte and icelandite to rhyolite</td><td>Afsp + ol + cpx + mt + ilm + ap ± aen ± qtz</td><td>Starkel, 2014; Henry et al., 2017</td></tr><tr><td>Green Tuff</td><td>continental rift</td><td>Cinque Denti caldera (Pantelleria, Italy)</td><td>45–47 ka</td><td>10 km³</td><td>pantellerite to comenditic trachyte</td><td>anorth + amph + mt + ilm + cpx + bt ± pl ± ol ± chev</td><td>Williams (2010); Williams et al. (2014); Liszewska et al. (2018)</td></tr><tr><td>Ignimbrite 'A'</td><td>oceanic hotspot</td><td>Tejeda caldera (Gran Canaria, Spain)</td><td>13.63 Ma</td><td>30 km³</td><td>phonolite</td><td>san + pl + bt + cpx + amph + haü + tit + mt + ap</td><td>Troll and Schmincke (2002); Wolff et al. (2020)</td></tr><tr><td>Granadilla Tuff</td><td>oceanic hotspot</td><td>Las Canadas caldera (Tenerife, Canary Islands, Spain)</td><td>600 ka</td><td>35 km³</td><td>phonolite</td><td>Afsp + bt + cpx + mt + ilm ± haü ± pl ± ap ± tit ± amph ± ol</td><td>Ellis et al. (2022a); Troll and Schmincke (2002); Wolff et al. (2020); Edgar et al. (2002); this study</td></tr><tr><td>Poris Tuff</td><td>oceanic hotspot</td><td>Las Canadas caldera (Tenerife, Canary Islands, Spain)</td><td>273 ka</td><td>3.5 km³</td><td>phonolite</td><td>Anorth + haü + cpx + bt + tit + mt + ilm ± pl ± amph ± ol</td><td>Wolff et al. (2020); Edgar et al. (2002)</td></tr><tr><td>Abrigo Tuff</td><td>oceanic hotspot</td><td>Las Canadas caldera (Tenerife, Canary Islands, Spain)</td><td>169 ka<br>40 to &lt;20 ka and 6.4 ka</td><td>20 km³</td><td>phonolite</td><td>pl + cpx + opx + bt + amph + ol + ilm + mt + pl + cpx + opx + bt + amph + ol + ilm + mt + pl + san + qtz + bt + amph + opx + mt + ilm + zr + ap</td><td>Wolff et al. (2020); Edgar et al. (2002)</td></tr><tr><td>Cuicultic Tuff</td><td>subduction</td><td>Los Humeros caldera (Mexico)</td><td></td><td>0.3 km³</td><td>basaltic andesite to trachydacite</td><td>pl + cpx + opx + bt + amph + ol + ilm + mt + pl + cpx + opx + bt + amph + ol + ilm + mt + pl + san + qtz + bt + amph + opx + mt + ilm + zr + ap</td><td>This study</td></tr><tr><td>Xoxoctic Tuff</td><td>subduction</td><td>Los Humeros caldera (Mexico)</td><td>50.7</td><td>0.6 km³</td><td>andesite to trachydacite</td><td>san + pl + bt + cpx + amph + mt + ilm</td><td>This study</td></tr><tr><td>Caetano tuff</td><td>extensional</td><td>Caetano caldera (Nevada, USA)</td><td>33.8 Ma</td><td>km³</td><td>rhyolitic</td><td>qtz + pl + san + cpx + opx + bt + amph + ilm + mt</td><td>Ellis et al. (2022a); this study</td></tr><tr><td>Pepom tephras</td><td>oceanic hotspot</td><td>Sete Cidades volcano (São Miguel, Azores)<br>Central Snake River plain volcanic field (northwestern Nevada, USA)</td><td>4.6 ka</td><td>DRE</td><td>trachytic</td><td>pl + cpx + opx + bt + amph + mt + ilm + zr + ap</td><td>Ellis et al. (2022b); this study</td></tr><tr><td>Jarbridge Rhyolite</td><td>extensional</td><td></td><td>16.09 to 15.09 Ma</td><td>7000 km³</td><td>rhyolitic</td><td>pl + cpx + opx + bt + amph + ilm + mt</td><td>Brueseke et al. (2014); this study</td></tr></tbody></table>
lava (118.7 ± 10.0 ka) while the other major rhyolitic eruptions occurred later with the Nikia lava (73.3 ± 5.8 ka) (ages are taken from Popa et al., 2020). Data from the KNVC used here is from Ellis et al. (2022a) in the case of the KPT, with additional analyses of KPT, Avlakli (12379), and Nikia (G10, N39) acquired during this study.

### 2.1.2. Carpenter Ridge Tuff - Southern Rocky Mountain volcanic field, Colorado, USA

The Carpenter Ridge Tuff (~27.55 Ma; CRT; Lipman and McIntosh, 2008) is a voluminous ignimbrite that is chemically zoned with early-erupted crystal-poor rhyolite (5–10 vol% crystals) transitioning to dacitic-trachytic crystal-rich (15–40 vol% crystals) deposits (Dorais et al., 1991; Bachmann et al., 2014). The mineral assemblage consists of sanidine, plagioclase, biotite, clinopyroxene, amphibole, magnetite, ilmenite and zircon. Data used in this study are taken from the compilation of Wolff et al. (2020).

### 2.1.3. Los Humeros caldera - Cuicuiltic Tuff and Xoxoctic Tuff - Mexico

The Los Humeros caldera is located in a subduction zone at the eastern part of the Mexican Volcanic Belt (Ferrari et al., 1999). The deposits for this study are the Xoxoctic Member (XM; 50.7 ± 4.4 ka; Carrasco-Núñez et al., 2018) and pumice-fall layers that alternated with scoriae layers that are called Cuicuiltic Member (CM; 40 to <20 ka and 6.4 ka; Ferriz and Mahood, 1984; Dávila-Harris and Carrasco-Núñez, 2014). The CM represents the last explosive event at Los Humeros caldera (Ferriz and Mahood, 1984) with a composition ranging from basaltic andesite to trachydacite (Dávila-Harris and Carrasco-Núñez, 2014). The compositional diversity of the CM is displayed by a range of mineral phases occurring in its end-members and including plagioclase, two pyroxenes, biotite, hornblende, olivine and Fe-Ti oxides. The Xoxoctic Tuff has a mineral assemblage dominated by olivine, plagioclase, clinopyroxene and Fe-Ti oxides.

## 2.2. Extensional Settings

### 2.2.1. Bishop Tuff - California, USA

The caldera-forming Bishop Tuff (BT; 0.76 Ma) erupted from the Long Valley magma chamber. It started as a Plinian eruption in the south-central part of what later became the caldera (Hildreth and Mahood, 1986). The BT consists of widespread fall deposits and non-welded to welded ignimbrites (Wilson and Hildreth, 1997). It is compositionally stratified and becomes less silicic with an increase in phenocryst contents upwards with ignimbrite 1(Ig1) having up to 10 wt % and ignimbrite 2 (Ig2) up to 10–24 wt% crystal contents (Hildreth, 1979; Hildreth, 1981; Hildreth and Wilson, 2007). The mineral assemblage consists of sanidine, quartz, plagioclase, magnetite, biotite, ilmenite, zircon and apatite (Hildreth, 1979). Data used in this paper are taken from Ellis et al. (2022a). Samples described in Ellis et al. (2022a) were further analysed in this study.

### 2.2.2. Green Tuff - Pantelleria, Italy

The Green Tuff (GT; 45.7 ± 1 ka; Scaillet et al., 2013) is the product of a Plinian caldera-forming eruption. The GT is comprised of Plinian fall deposit overlain by a strongly welded ignimbrite (Williams et al., 2014) that is chemically zoned from pantellerite (base) to trachyte (top) with pantellerite being the dominant magma type. The mineral assemblage of the GT consists of alkali feldspar, olivine, clinopyroxene, ilmenite, apatite, aenigmatite and quartz. Data used in this study come from the compilation of Wolff et al. (2020).

### 2.2.3. Caetano Tuff - North-central Nevada, USA

The Caetano Tuff (CT; ~33.8 Ma; John et al., 2008) was a voluminous caldera-forming eruption (e.g., Lipman et al., 1972; Christiansen and Yeats, 1992), consisting of two ash-flow units with the older, more mafic outflow tuff (ca. 34.2 Ma), and the main caldera-filling Caetano Tuff (ca. 33.8 Ma) that is related to the outflow tuff (Stewart and McKee,

1977). The upper unit is a crystal-rich, rhyolitic ash-flow tuff (John et al., 2008) with a mineral assemblage consisting of plagioclase, sanidine, quartz, biotite, hornblende, orthopyroxene, magnetite, ilmenite, zircon and apatite (Watts et al., 2016). The samples selected for this study are from a calc-alkaline suite of the crystal-rich rhyolite portion with one representative of the glassy portion and the other collected from the devitrified part. Data used in this study are taken from Ellis et al. (2022a) and additional analyses of samples from their study were collected here.

### 2.2.4. Jarbidge Rhyolite - North-eastern Nevada, USA

The Jarbidge Rhyolite (16.09 ± 0.08 Ma to 15.09 ± 0.06 Ma; Brueseke et al., 2014) is a series of predominantly effusive, crystal-rich (c. 40 vol%) lavas. Jarbidge Rhyolite is found at approximately the same location, but a few million years prior to the main stage of rhyolite from the Snake River Plain volcanic province (Perkins and Nash, 2002) and the debate about the extent to which it belongs to the Yellowstone province is ongoing. Jarbidge Rhyolite has a mineral assemblage indicative of xenocryst entrainment including quartz, plagioclase, sanidine, clinopyroxene, orthopyroxene, biotite and amphibole (often altered) and Fe-Ti oxides, apatite and zircon (Brueseke et al., 2014). While entirely rhyolitic in bulk composition, the Jarbidge Rhyolite notably exists as high-Ba and low-Ba varieties. In this study we investigated both types (low-Ba: JC09–5, JC08–10, JC09–16; higher-Ba: JC09–16).

## 2.3. Subduction-extensional settings

### 2.3.1. Campi Flegrei caldera - Campanian Plain - Italy

The active Campi Flegrei caldera (Campanian Plain, southern Italy; e.g., Scandone and Patacca, 1984) has had two explosive caldera-forming events, (1) Campanian Ignimbrite (CI; ~39 ka, De Vivo et al., 2001), and (2) Neapolitan Yellow Tuff (NYT; 14.9 ka; Deino et al., 2004).

The CI is a voluminous and widespread pyroclastic sequence composed of a basal fallout deposit and distinct pyroclastic flow units (Civetta et al., 1997). The composition of the CI ranges from a relatively crystal-poor (5–6 vol%) phonolite at the base to a crystal-rich (~36 vol%) trachyte on passing up-succession (Forni et al., 2016). The mineral assemblage consists of feldspars, clinopyroxenes, biotites, oxides and accessory apatite (Civetta et al., 1997).

The Astroni volcano produced seven small eruptions (0.5 km³; Isaia et al., 2004) between 4.8 and 3.8 ka (Isaia et al., 2004) with compositions including tephrifronolites, phonolites and trachytes, and a mineralogy typified by feldspar, clinopyroxene, biotite, Fe-Ti oxides and apatite (Isaia et al., 2004). Data used in this paper are taken from Forni et al. (2016, 2018) and Ellis et al. (2022a) with samples petrographically described in Ellis et al. (2022a) taken for additional analyses in this study.

### 2.3.2. Peach Spring Tuff - Arizona-California-Nevada, USA

The Peach Spring Tuff (PST; 18.78 ± 0.02 Ma; Ferguson et al., 2013) is a predominantly welded pyroclastic deposit. The compositions of the PST ignimbrite are trachyte, low-silica rhyolite, and high-silica rhyolite (Foley et al., 2020). The high-SiO₂ rhyolites erupted early and are crystal-poor (5–10 vol%) whereas low-SiO₂ rhyolites and trachytes erupted later and are crystal-rich (20–23 vol%; Pamukcu et al., 2013; Foley et al., 2020). The mineral assemblage consists of sanidine, plagioclase, biotite, quartz, hornblende, pyroxene, titanite, ilmenite, magnetite, sphene, zircon, chevkinite and apatite. Data used in this study come from the compilation of Wolff et al. (2020) and Foley et al. (2020).

## 2.4. Hotspot settings

### 2.4.1. McDermitt Tuff - Nevada-Oregon, USA

The McDermitt volcanic field is a caldera complex that is associated
with some of the earliest silicic volcanism of the Columbia River–Yellowstone magmatic province (Coble and Mahood, 2012). Erup- tion of the McDermitt Tuff (MDT; 16.39 ± 0.02 Ma; Henry et al., 2017) formed the McDermitt caldera and the tuff is chemically zoned from high-silica rhyolite to trachydacite and Fe-rich andesite (icelandite) (Henry et al., 2017). The mineral assemblage of the MDT consists of anorthoclase, clinopyroxene, fayalite, amphibole, magnetite and ilmenite. Data used in this study come from the compilation of Wolff et al. (2020).

### 2.4.2. Ignimbrite 'A' - Gran Canaria, Spain

The ocean island volcano of Gran Canaria (Canary Islands, Spain) records >14 million years of submarine and subaerial volcanism from the Canarian hotspot. Explosive silicic volcanism started almost immediately upon cessation of the shield-forming stage and persisted for some 5 Myr prior to a volcanic hiatus (van den Bogaard and Schmincke, 1998). Within this period of silicic activity, the first phase was silica-saturated forming the Mogan Group prior to the eruption of the silica-undersaturated Fataga Group. Ignimbrite 'A' (13.63 ± 0.3 Ma; van den Bogaard and Schmincke, 1998) forms the lowest unit of the Upper Mogan sequence. Ignimbrite 'A' comprises compositionally zoned rhyolitic (base) to trachytic (top) deposits. The crystal contents increase from the base to the top. The mineral assemblage of ignimbrite 'A' consists of oligoclase, anorthoclase, amphibole, clinopyroxene, phlogopite, chevkinite, olivine, magnetite and ilmenite. Data used in this study come from the compilation of Wolff et al. (2020).

### 2.4.3. Las Cañadas caldera – Granadilla-Poris-Abrigo – Tenerife, Canary Islands, Spain

The Las Cañadas caldera (Tenerife, Canary Islands, Spain) is at the summit of a large phonolitic shield volcano (Bryan et al., 2002). Following a shield-building stage, silicic activity began at c. 1.57 Ma and the samples from this study are from younger portions of this activity, the Guajara (0.85–0.57 Ma) and Diego Hernández (0.53–0.196 Ma) formations (Martí et al., 1994, 1997). In this work we focused on the Granadilla pumice and literature data from the Poris and the El Abrigo deposits.

The Granadilla eruption (600 ± 9 ka; Brown et al., 2003) was a caldera-forming phonolitic plinian pumice deposit (Booth, 1973). The Granadilla ignimbrite is considered as crystal-poor but contains sanidine, plagioclase, biotite, pyroxene, amphibole, haüyne, titanite, magnetite, ilmenite and apatite (Bryan et al., 2002). Data used in this paper are taken from Ellis et al. (2022a). Samples described in Ellis et al. (2022a) were further analysed in this study.

The Poris deposits is a white to grey dominantly phonolitic non-welded ignimbrite (273 ± 5 ka; Brown et al., 2003) with a thin, fine-grained pumice at its base. The Poris deposit mineral assemblage consists of sanidine, labradorite, andesine, biotite, pyroxene, amphibole, haüyne and titanomagnetite (Bryan et al., 2002). Data used in this study come from the compilation of Wolff et al. (2020).

The El Abrigo ignimbrite (169 ± 1 ka; Brown et al., 2003) is the youngest, explosive, caldera-forming phonolite pyroclastic eruption of Tenerife. It is a pale grey, variably lithic-rich non-welded ignimbrite (Pittari et al., 2006). The mineral assemblage consists of sanidine, anorthoclase, oligoclase, biotite, pyroxene amphibole, ilmenite, titanomagnetite, sphene, haüyne–sodalite, nepheline and apatite. (Bryan et al., 2002; Pittari et al., 2006). Data used in this study come from the compilation of Wolff et al. (2020).

### 2.4.4. Sete Cidades volcano - Pepom tephra - São Miguel, Azores, Portugal

The Sete Cidades volcanic centre extends back to at least ~200 ka (Moore, 1990) and is described by Queiroz et al. (2015) and Beier et al. (2006). The samples used here represent the youngest explosive volcanic activity from Sete Cidades, the Pepom tephras (Queiroz et al., 2008; Ellis et al., 2022b). There are 17 defined trachytic units from P1 (oldest) to P17 (youngest) and they are crystal-poor to crystal-moderate with a

mineral assemblage of sanidine, plagioclase, biotite, clinopyroxene, amphibole and Fe-Ti oxides. Data used in this paper are taken from Ellis et al. (2022b).

## 3. Analytical methods

Samples were crushed and mineral phases were handpicked from the cleaned, crushed and sieved rocks under a binocular microscope. Prior to in-situ geochemical analyses, groundmass glass, sanidine, plagioclase and clinopyroxene were embedded in epoxy mounts or thin sections were used.

### 3.1. Major element measurements

Mineral phases and groundmass glass compositions were determined using an electron probe micro-analyser (EPMA; JEOL JXA 8200 Superprobe) at the ETH Zurich. Using operating conditions of 15 kV acceleration voltage, and a 15 nA beam current with counting times of 20 s on the peaks and 10 s on the background for glass and a 20 nA beam current for sanidine and plagioclase. Groundmass glass analyses were measured with a defocused beam of 20 µm while sanidine and plagioclase were analysed with a 10 µm beam diameter in order to minimise alkali (Na, K) migration. To ensure reproducibility of <5% for all major and minor elements different mineral standards were used (see supplementary data).

### 3.2. Trace element measurements

Trace element contents of the mineral phases and the groundmass glass were determined by laser ablation inductively coupled plasma mass spectrometry (LA-ICP-MS) at the Institute of Geochemistry and Petrology, ETH Zurich. A 193 nm Resonetics ArF excimer laser ablation system was used, coupled to a Thermo Element XR sector-field ICP-MS. Details about the methods can be found in Neukampf et al. (2019). Typical laser ablation diameters were 43 µm for mineral phases whereas it was 19 µm for groundmass glass. For mineral phases and glass the NIST SRM610 or NIST SRM612 glass wafers were used as the primary reference material whereas GSD-1G basalt glass (USGS) was employed as a secondary reference material to ensure the accuracy and reproducibility of the instrumental approach. All data were reduced using SILLS (Guillong et al., 2008) and precision of the analyses is considered to be better than 10% relative as Li is almost always comfortably above detection limits.

### 3.3. Lithium isotope measurements

Mineral and groundmass glass separates were handpicked from crushed and sieved rocks under a binocular microscope. All Li concentration and isotope measurements of samples were performed using a Neptune multiple-collector inductively-coupled-plasma mass spectrometer (MC-ICP-MS), housed at the Czech Geological Survey. Details of the purification and analytical/instrumental procedures are given elsewhere (Magna et al. (2004, 2006a). Mineral separates (11–27 mg) were digested in a mixture of 15 M HNO₃ and 27 M HF (1/6 v/v) for 48 h on a hot plate (130 °C). The evaporated samples were then refluxed three times with small amounts of 15 M HNO₃ and finally equilibrated in 6 M HCl for 24 h on a hot plate (80 °C). Lithium was isolated from the mineral separates via a two-step cation exchange column procedure, first employing the BioRad AG50W-X8 (mesh 200–400) cation-exchange resin and 80% methanol–1 M HNO₃ mixture as elution medium, followed by additional cleaning using the BioRad AG50W-X12 (mesh 200–400) cation-exchange resin and 0.5 M HNO₃ as elution medium. We analysed aliquots from each purified sample and the standards for all elements using an Agilent 7900x single collector ICPMS, housed at the Czech Geological Survey, to investigate any potential impurities.

The lithium concentrations were analysed in clean Li fractions
extensional setting have Li concentrations ranging 6.9–80.0 ppm (n = 124; this study). The effusive samples have on average higher Li concentrations relative to explosive samples. Samples from extensional settings have the lowest Li concentrations, followed by samples from subduction zones, oceanic hotspot settings, subduction-extensional setting and with samples from a continental hotspot setting having the largest range (the ranges of Li contents are unrelated to the number of analyses).

## 4.2. Quartz

Lithium data (n = 283) for quartz are only available for several rhyolitic eruptions (Fig. 3; Table 3) from subduction zones and extensional settings. The Li concentrations in quartz crystals from the subduction setting range between 3.2 and 12.8 ppm (n = 203; KPT: rhyolitic sample: 3.2–4.9 ppm; this study; granitic samples: 5.2–12.8 ppm; Fiedrich et al., 2020) and in quartz grains from the extensional settings from 6.1 to 21.1 ppm (n = 80; Caetano Tuff: glassy sample: 16.3–21.1 ppm; devitrified sample: 14.6–20.6 ppm). On the isoatomic Al:Li 1:1 line (Fig. 3), which indicates if the Li concentrations are sufficiently high to charge balance Al, only the samples from extensional settings plot around the 1:1 line (this is further discussed below). Quartz crystals from the BT show a small increase in Li concentrations towards the rim which is associated with hydrogen as determined by Jollands et al. (2020).

## 4.3. Plagioclase

In total, 114 data points are available for plagioclase from the alkaline samples (Fig. 4 and Table 3). Plagioclase crystals in the explosive calc-alkaline samples display a Li range from 1.4 to 19.7 ppm (n = 374). The plagioclase crystals from the subduction-extensional setting have Li ranges from 0.9 to 25.1 ppm (Forni et al., 2016, 2018). The Li contents for samples from the subduction zone settings have Li concentrations between 1.4 and 19.4 ppm (n = 174) and for the extensional setting from 9.1 to 19.7 ppm (n = 200).

In the effusive calc-alkaline samples from a subduction zone setting the Li concentrations in plagioclase crystals vary from 5.0 to 35.0 ppm (n = 136; Fig. 4; Table 4) and plagioclase grains from an extensional setting have Li contents from 4.8 to 30.8 ppm (n = 172).

<figure><img src="image_7.png" /><figcaption>Fig. 3. Plot of Li vs. Al concentrations in quartz from different units displayed by their tectonic settings (subduction zone and extensional settings) and cooling regimes (devitrified representing slower cooled samples and glassy representing quicker cooled samples) The isoatomic Al:Li 1:1 line is given for reference and illustrates the ideal Si⁴⁺ ↔ Al³⁺ + Li⁺ substitution.</figcaption></figure>

<table><caption>**Table 3**

Li concentrations (in ppm) in quartz crystals from different deposits.</caption><thead><tr><th>Deposit</th><th>min</th><th>max</th><th>avg.</th><th>n</th><th>Reference</th></tr></thead><tbody><tr><td>Kos Plateau Tuff (granitic sample)</td><td>5.2</td><td>12.8</td><td>8.9 ± 0.3</td><td>131</td><td>Fiedrich et al., 2020</td></tr><tr><td>Kos Plateau Tuff (rhyolitic sample)</td><td>3.2</td><td>4.9</td><td>3.9 ± 0.8</td><td>72</td><td>Fiedrich et al., 2020</td></tr><tr><td>Bishop Tuff</td><td>6.1</td><td>13.4</td><td>2.0 ± 17.5</td><td>55</td><td>This study</td></tr><tr><td>Caetano Tuff (devitrified sample)</td><td>14.6</td><td>20.6</td><td>1.7 ± 19.0</td><td>10</td><td>This study</td></tr><tr><td>Caetano Tuff (glassy sample)</td><td>16.3</td><td>21.1</td><td>1.3 ± 1.3</td><td>15</td><td>This study</td></tr></tbody></table>

<figure><img src="image_8.png" /><figcaption>Fig. 4. Plot of Ba vs. Li compositions in plagioclase crystals of different eruption units sorted by their tectonic settings (subduction-extensional setting as well as subduction zones and extensional settings). Barium is used as it reflects the melting and subsequent crystallisation of sanidine from melts produced from a feldspar-dominated cumulate.</figcaption></figure>

<table><caption>**Table 4**

Li concentrations (in ppm) in plagioclase crystals from different deposits.</caption><thead><tr><th>Deposit</th><th>min</th><th>max</th><th>avg.</th><th>n</th><th>Reference</th></tr></thead><tbody><tr><td>alkaline trend</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>Campi Flegrei calc-alkaline trend</td><td>0.9</td><td>25</td><td>8.8 ± 5.4</td><td>114</td><td>Forni et al. (2016, 2018)</td></tr><tr><td>Bishop Tuff</td><td>9.1</td><td>19.7</td><td>13.3 ± 2.7</td><td>200</td><td>This study</td></tr><tr><td>Kos Plateau Tuff (granitic sample)</td><td>10</td><td>11.7</td><td>10.9 ± 0.5</td><td>17</td><td>This study</td></tr><tr><td>Xoxtic Tuff</td><td>1.4</td><td>19.4</td><td>3.5 ± 10.0</td><td>117</td><td>This study</td></tr><tr><td>Ciutilitic Tuff calc-alkaline lavas</td><td>7.9</td><td>13.8</td><td>1.4 ± 31.8</td><td>40</td><td>This study</td></tr><tr><td>Nikia flow (G10)</td><td>27</td><td>35</td><td>1.9 ± 15.8</td><td>50</td><td>This study</td></tr><tr><td>Nikia flow (N39)</td><td>5.0</td><td>23</td><td>4.4 ± 23.2</td><td>36</td><td>This study</td></tr><tr><td>Avlaki lava flow (12379)</td><td>18</td><td>32</td><td>3.5 ± 8.9</td><td>50</td><td>This study</td></tr><tr><td>high-Ba Jarbridge Rhyolite</td><td>4.8</td><td>17.9</td><td>3.3 ± 17.3</td><td>29</td><td>This study</td></tr><tr><td>low-Ba Jarbridge Rhyolite</td><td>5.1</td><td>30.8</td><td>5.3 ± 5.3</td><td>143</td><td>This study</td></tr></tbody></table>
## 4.4. Sanidine

In the alkaline samples there are 944 analyses of sanidine crystals with Li contents between 0.5 and 38.7 ppm (Fig. 5a; Table 5). The plagioclase crystals from an oceanic hotspot setting have the lowest Li concentrations ranging from 0.5 to 3.4 ppm (n = 315) and for the samples from the subduction-extension setting Li contents vary between 0.9 and 38.7 (n = 629). The large variations in Ba in the subduction-extensional setting reflect the melting and subsequent crystallisation of sanidine from melts produced from a feldspar-dominated cumulate. The explosive calc-alkaline samples (Fig. 5b) yielded 122 analyses in total. The sanidine grains from the subduction zones display Li contents between 3.5 and 6.0 ppm (n = 71) and samples from extensional settings have Li concentrations between 0.3 and 6.1 ppm (n = 51; Caetano Tuff devitrified sample: 0.3–6.1 ppm; glassy sample: 1.2–2.3 ppm).

The effusive calc-alkaline samples from an extensional setting have Li concentrations between 0.9 and 8.1 ppm (n = 267). The sanidine crystals of samples from hotspot and subduction-extensional settings have higher Li concentrations compared to samples from subduction and extensional settings.

## 4.5. Biotite

In the alkaline samples, a total of 1012 analyses of biotite crystals are available (Fig. 6a; Table 6). The biotites present in the samples from the hotspot setting have Li concentrations between 0.8 ppm and 86.6 ppm (n = 367; Ellis et al., 2022a, 2022b) and biotite crystals found in samples from subduction-extensional settings have Li concentrations ranging from 4.7 to 99.3 ppm (n = 645; Forni et al., 2018; Ellis et al., 2022a). Lithium concentrations in biotite crystals from an extensional setting ranging from 10.2 to 42.2 ppm (Caetano devitrified: 10.2–42.2 ppm, n = 21; glassy part: 11.9–34.0 ppm; n = 27; Ellis et al., 2022a).

We do not present the Li concentrations data for biotite from BT and KPT. Ellis et al. (2022a) have shown that the Li concentrations measured in these biotite crystals mostly are linked to fluid inclusions hosted in the interlayer of biotite and may thus not represent the Li concentrations of the biotite crystals. In calc-alkaline systems biotite often is a late stage phase and in a number of systems it is entirely absent (e.g., Mesa Falls Tuff).

## 4.6. Pyroxene and olivine

The Li concentrations in pyroxene crystals from the alkaline samples are 0.6 to 31.0 ppm (n = 1558; Forni et al., 2016, 2018; Ellis et al.,

<table><caption>**Table 5**

Li concentrations (in ppm) in sanidine crystals from different deposits.</caption><thead><tr><th>Deposit</th><th>min</th><th>max</th><th>avg.</th><th>n</th><th>Reference</th></tr></thead><tbody><tr><td>alkaline trend</td><td></td><td></td><td>3.7 ±</td><td></td><td></td></tr><tr><td>Campi Flegrei</td><td>0.9</td><td>38.7</td><td>4.7</td><td>629</td><td>Forni et al. (2016, 2018)</td></tr><tr><td>Pepom tephras</td><td>1.4</td><td>19.4</td><td>1.5 ±<br>0.5</td><td>285</td><td>Ellis et al. (2022b)</td></tr><tr><td>Granadilla Tuff<br>calc-alkaline trend</td><td>1.0</td><td>1.4</td><td>1.2 ±<br>0.1</td><td>30</td><td>This study</td></tr><tr><td>Bishop Tuff</td><td>3.6</td><td>6.0</td><td>4.6 ±<br>0.5</td><td>56</td><td>This study</td></tr><tr><td>Kos Plateau Tuff (granitic sample)</td><td>3.5</td><td>6.0</td><td>4.5 ±<br>0.5</td><td>71</td><td>This study</td></tr><tr><td>Caetano Tuff (devitrified sample)</td><td>0.3</td><td>6.1</td><td>1.4 ±<br>1.5</td><td>34</td><td>This study</td></tr><tr><td>Caetano Tuff (glassy sample)</td><td>1.2</td><td>2.3</td><td>1.6 ±<br>0.3</td><td>16</td><td>This study</td></tr><tr><td>calc-alkaline lavas</td><td></td><td></td><td>2.0 ±<br>0.4<br>3.8 ±</td><td></td><td></td></tr><tr><td>high-Ba Jarbidge Rhyolite</td><td>1.2</td><td>3</td><td></td><td>129</td><td>This study</td></tr><tr><td>low-Ba Jarbidge Rhyolite</td><td>0.9</td><td>8.1</td><td>1.7</td><td>138</td><td>This study</td></tr></tbody></table>

<figure><img src="image_9.png" /><figcaption>**Fig. 6.** Plot of Nb vs. Li concentrations in biotite crystals from oceanic hotspots and a subduction-extensional setting with alkaline compositions. Here Nb is used as it is considered to be incompatible.</figcaption></figure>

<figure><img src="image_10.png" /><figcaption>**Fig. 5.** Plot of Ba vs. Li concentrations in sanidine crystals of different eruption units organised by their tectonic settings. Barium is used as it reflects the melting and subsequent crystallisation of sanidine from melts produced from a feldspar-dominated cumulate. (A) Plot for sanidine crystals from alkaline units (from oceanic hot spots and the subduction-extensional settings) with the grey field indicating the range of the sanidine compositions shown in (B). (B) Plot for sanidine crystals from calc-alkaline units from subduction zones and extensional tectonic settings.</figcaption></figure>
<table><caption>**Table 6**
Li concentrations (in ppm) in biotite crystals from different deposits.</caption><thead><tr><th>Deposit</th><th>min</th><th>max</th><th>avg.</th><th>n</th><th>Reference</th></tr></thead><tbody><tr><td>alkaline trend</td><td></td><td></td><td>19.8 ±</td><td></td><td>Forni et al. (2016, 2018); Ellis et al. (2022a)</td></tr><tr><td>Campi Flegrei</td><td>4.7</td><td>99.3</td><td>16.4<br>12.8 ±</td><td>326</td><td></td></tr><tr><td>Pepom tephras</td><td>0.8</td><td>86.6</td><td>14.0<br>18.7 ±</td><td>645</td><td>Ellis et al. (2022a)</td></tr><tr><td>Granadilla tuff<br>calc-alkaline trend</td><td>11.2</td><td>23.2</td><td>2.6</td><td>41</td><td>Ellis et al. (2022b)</td></tr><tr><td>Caetano Tuff<br>(devitrified)</td><td>10.2</td><td>42.2</td><td>20.5 ±<br>8.7</td><td>21</td><td>Ellis et al. (2022a)</td></tr><tr><td>Caetano Tuff<br>(glassy)</td><td>11.9</td><td>34</td><td>19.9 ±<br>5.8</td><td>27</td><td>Ellis et al. (2022a)</td></tr></tbody></table>

2022b) with pyroxene grains from an ocean island setting having Li contents between 2.2 and 19.8 ppm (n = 336; Ellis et al., 2022b) and between 0.6 and 31.0 ppm for samples from a subduction-extensional setting (n = 822). In the calc-alkaline samples from a subduction zone setting the Li concentrations in pyroxene grains are between 0.7 and 21.8 ppm (n = 93; Fig. 7b, Table 7) and in olivine crystals between 1.0 and 16.3 ppm (n = 96). Therefore, pyroxene crystals from the oceanic hotspot settings as well as from samples from the subduction-extensional setting have higher Li concentrations compared to pyroxene grains from subduction zone samples.

## 4.7. Lithium isotope systematics

The lithium isotopic compositions of mineral separates and bulk samples show a large variability exceeding 10‰ (Fig. 8). The isotopically heaviest samples are the quartz from the rhyolitic samples from the extensional setting (Caetano Tuff and Bishop Tuff) ranging from 10.0 to 11.5‰ whereas quartz from the subduction zone sample (Kos Plateau Tuff) is isotopically lighter and varies between 2.0 and 2.4‰.

In samples from the subduction-extensional setting the $\delta^{7}\text{Li}$ values of the phonolitic groundmass glass range from 0.5‰ to 0.9‰ (Astroni and Campanian ignimbrite, respectively), clinopyroxene values vary from 0.5‰ to -0.3‰ (Astroni and Campanian ignimbrite, respectively) whereas the sanidine crystals show negative $\delta^{7}\text{Li}$ values (-0.9‰ and -5.1‰, Astroni and Campanian ignimbrite, respectively). The lithium isotopic compositions of bulk rock and biotite separates of different samples, for which we collected new Li data in this study, are published in Ellis et al. (2022a).

<figure><img src="image_11.png" /><figcaption>**Fig. 7.** Plot of Mg vs. Li concentrations in pyroxene and olivine crystals organised by their tectonic settings (A: oceanic hot spots and subduction-extensional settings; B: subduction zone settings). Magnesium is used as Li usually enters the octahedrally co-ordinated Mg site in olivine and pyroxene due to their almost identical ionic radii. (A) Plot for pyroxene from alkaline units. (B) Plot for clinopyroxene, orthopyroxene and olivine (square symbol) from calc-alkaline units with the grey field indicating the range of the sanidine compositions shown in (A).</figcaption></figure>

<table><caption>**Table 7**
Li concentrations (in ppm) in clinopyroxene (cpx), orthopyroxene (opx) and olivine (ol) from different deposits.</caption><thead><tr><th>Deposit</th><th>min</th><th>max</th><th>avg.</th><th>n</th><th>Reference</th></tr></thead><tbody><tr><td>Campi Flegrei (cpx)</td><td>0.6</td><td>31</td><td>4.8<br>4.8 ±</td><td>822</td><td>Forni et al. (2016, 2018)</td></tr><tr><td>Pepom tephras (cpx)</td><td>2.2</td><td>19.8</td><td>2.0<br>15.8 ±</td><td>336</td><td>Ellis et al. (2022b)</td></tr><tr><td>Xoxoctic Tuff (Layer A-1; cpx)</td><td>12</td><td>21.8</td><td>3.7<br>1.4 ±</td><td>6</td><td>This study</td></tr><tr><td>Xoxoctic Tuff (Layer A-2, B, D; cpx)</td><td>0.7</td><td>2.6</td><td>0.4<br>3.0 ±</td><td>81</td><td>This study</td></tr><tr><td>Xoxoctic Tuff (Layer D; opx)</td><td>2.6</td><td>3.2</td><td>0.2<br>13.0 ±</td><td>6</td><td>This study</td></tr><tr><td>Xoxoctic Tuff (Layer A-1; ol)</td><td>9.7</td><td>16.3</td><td>3.3<br>2.9 ±</td><td>2</td><td>This study</td></tr><tr><td>Xoxoctic Tuff (Layer A-2, B, D; ol)</td><td>1.0</td><td>5.7</td><td>0.9<br>0.9 ±</td><td>94</td><td>This study</td></tr></tbody></table>

# 5. Discussion

Based on the combination of new analyses from this study and literature data we are now able to constrain, for the first time, the Li abundance for mineral phases and melt compositions ranging from basaltic trachy-andesite to phonolite/trachyte (Table 8) and from andesite to rhyolite which are summarised in Table 2 and Fig. 9. The first-order observations made in this study, combined with published data, make a convincing case that the Li concentrations in the groundmass glass and minerals are clearly independent of the tectonic setting, and that there is no correlation with the age of the eruption or the style of volcanic effusive activity (fallout versus a quickly cooled non-welded lava). With a large amount of data, we can start to identify outliers that may have experienced significant post-eruptive modification and such data can help to potentially identify whether such modifications are more prevalent in composition or regimes of specific eruptive dynamics.

## 5.1. Partition coefficients

Most work done to address the crystal/melt partitioning of Li has focused on less evolved, mafic magmatic systems (e.g., Blundy et al., 1998; Brenan et al., 1998a, 1998b; Ottolini et al., 2009). More evolved systems require a significantly more varied approach (e.g., Iveson et al., 2018; Neukampf et al., 2019) due to a much wider compositional range of melts with >52 wt% SiO$_2$, resulting in a significantly greater mineralogical variability. Partition coefficients for basaltic systems have been
<figure><img src="image_12.png" /><figcaption>**Fig. 8.** Plots of δ⁷Li for different mineral phases and bulk-rock separates with Li concentrations provided next to the symbols. Plots are organised by cooling environment (quenched, vitrophyre and slowly cooled). Quickly cooled deposits are represented by Mesa Falls Tuff (MFT), Kos Plateau Tuff (KPT), Bishop Tuff (BT), Astroni volcano (Astroni) and Granadilla eruption (GP). Larger symbols for Mesa Falls Tuff (MFT) represent δ⁷Li for mineral and glass separates (plagioclase and groundmass glass) measured by MC-ICP-MS. Smaller symbols for the plagioclase and groundmass glass of MFT represent SIMS measurements (Li concentrations given in italic). All other symbols in the figure represent samples measured by MC-ICP-MS. The δ⁷Li for biotite separates from BT and KPT are excluded from this plot as they are related to another process (see Ellis et al., 2022a). The rapidly cooled vitrophyre is represented by samples of the Tuff of Knob (Knob), Caetano Tuff (Caetano) and Campanian Ignimbrite (Campanian) and the slowly cooled ignimbrite is represented by samples of the Tuff of Knob and the Caetano Tuff. Some of the data is taken from Ellis et al. (2018, 2022a).</figcaption></figure>
<table><caption>**Table 8**
Li concentrations (in ppm) for the different mineral phases and co-existing glasses.</caption><thead><tr><th rowspan="2"></th><th colspan="3">alkaline trend</th><th colspan="3">calc-alkaline trend</th></tr><tr><th>min</th><th>max</th><th>average</th><th>min</th><th>max</th><th>average</th></tr></thead><tbody><tr><td>glass</td><td>1.1</td><td>318</td><td>26.2</td><td>1301</td><td>176</td><td>33.8</td><td>2115</td></tr><tr><td>biotite</td><td>0.8</td><td>99.3</td><td>15.4</td><td>1012</td><td>2314</td><td>412</td><td>218</td></tr><tr><td>plagioclase</td><td>0.9</td><td>25.1</td><td>8.8</td><td>114</td><td>35.0</td><td>14.5</td><td>682</td></tr><tr><td>clinopyroxene</td><td>0.6</td><td>31.0</td><td>5.6</td><td>1158</td><td>21.8</td><td>2.6</td><td>93</td></tr><tr><td>sanidine</td><td>0.5</td><td>38.7</td><td>3</td><td>920</td><td>8.1</td><td>3.2</td><td>494</td></tr><tr><td>quartz</td><td>-</td><td>-</td><td>-</td><td>-</td><td>21.1</td><td>8.5</td><td>278</td></tr><tr><td>olivine</td><td>-</td><td>-</td><td>-</td><td>-</td><td>16.3</td><td>3.3</td><td>96</td></tr></tbody></table>

<table><caption>**Table 9**
Apparent partition coefficients for Li in the co-existing mineral phases.</caption><thead><tr><th rowspan="2"></th><th colspan="3">alkaline trend</th><th colspan="3">calc-alkaline trend</th></tr><tr><th>min</th><th>max</th><th>n</th><th>min</th><th>max</th><th>n</th></tr></thead><tbody><tr><td>biotite</td><td>0.01</td><td>2.65</td><td>549</td><td>0.53</td><td>47.54</td><td>98</td></tr><tr><td>plagioclase</td><td>0.01</td><td>0.29</td><td>53</td><td>0.06</td><td>0.81</td><td>245</td></tr><tr><td>clinopyroxene</td><td>0.01</td><td>0.48</td><td>485</td><td>0.03</td><td>0.99</td><td>39</td></tr><tr><td>sanidine</td><td>0.01</td><td>0.24</td><td>463</td><td>0.02</td><td>0.15</td><td>149</td></tr><tr><td>quartz</td><td>-</td><td>-</td><td>-</td><td>0.09</td><td>0.94</td><td>92</td></tr><tr><td>olivine</td><td>-</td><td>-</td><td>-</td><td>0.04</td><td>0.44</td><td>45</td></tr></tbody></table>

determined experimentally and are summarised elsewhere (Tomascak et al., 2016; Penniston-Dorland et al., 2017). Although large ranges in Kds are reported for Li, we note that even some typical trace elements in mineral phases, such as Sr in plagioclase, display ranges of Kd values (0.9–33 from the GERM database). In a single experimental study, Ice-nhower and London (1996) obtained a range of Kds from 1.1 to 18.4 for Ba in sanidine. We note that these ranges are for simple trace elements that occupy one site in the crystal lattice. Conversely, Li has the potential to occupy both the lattice site and interstitial site. This makes it significantly more complicated to characterise in terms of more accurately known Kd values.

With the combination of literature data and new trace element abundances for different mineral phases and co-existing groundmass glass in alkaline and silica-rich systems we are able to calculate apparent partition coefficients ($K_{Li(min/melt)}$, Fig. 10; see Supplement data 2) for the different deposits. Here, we only use the glass with the highest Li concentration as we assume that this represents the least degassed and least affected by potential hydration values which is a different approach compared to that taken by Neukampf et al. (2019). Melt inclusion data were not used for the calculation as they most likely are not in equilibrium with the host mineral phases because the Li concentration of the mineral phases could have been severely modified by different processes (e.g., degassing; Neukampf et al., 2021; Neukampf et al., 2022). For the mineral phases considered in this study we used the Li composition from the rims because they are more likely to be in equilibrium with the surrounding glass. All rim data of the minerals are used without filtering. Our results confirm that for more evolved alkaline system Li is a moderately incompatible element in all coexisting phases,

which has previously been shown for basaltic systems (olivine: Kd = 0.20–0.56; plagioclase: Kd = 0.15–0.7;), clinopyroxene: Kd = 0.11–0.59; orthopyroxene: Kd = 0.17–0.26; Ryan, 1989; Hart and Dunn, 1993; Blundy et al., 1998; Brenan et al., 1998a; Taura et al., 1998; McDade et al., 2003; Aigner-Torres et al., 2007; Ottolini et al., 2009; Beard et al., 2019). Likewise, for calc-alkaline systems it has previously been determined that Li is moderately incompatible in all mineral phases like fayalite (Kd = 0.39 ± 0.09), plagioclase (Kd = 0.23 ± 0.02), quartz (Kd = 0.45 ± 0.8), clinopyroxene (Kd = 0.26 ± 0.04), orthopyroxene (Kd = 0.20 ± 0.02) and sanidine (Kd = 0.10 ± 0.08) (Iveson et al., 2018; Neukampf et al., 2019).

### 5.1.1. Apparent partition coefficients for mafic mineral phases

In the studied alkaline systems, biotite most often has the highest apparent partition coefficient with biotite from the hotspot settings (Granadilla Tuff, Pepom tephra) displaying the highest values which range from 0.17 to 2.65. Biotite grains from a subduction-extensional setting (Campi Flegrei) yield the lowest Kds with a mean value of 0.11 (CI: 0.01–0.30; Astroni: 0.24–0.44; NYT: 0.05–0.35). This shows that in alkaline systems, Li is moderately incompatible in biotite (except for biotite from the hotspot setting), similar to biotite from the calc-alkaline trend (from an extensional setting: Caetano Tuff: 0.53–1.51). The Kds for the biotite crystals from the alkaline hotspot settings and the from the calc-alkaline extensional setting are very similar, whereas Kds from a subduction-extensional setting are lower in comparison. Based on the Kds it impossible to distinguish between the influence of the tectonic setting and the chemical trends. The Kds for the other calc-alkaline biotites are not considered here because, as stated in Section 4.5, the elevated Li concentrations of these biotites are due to fluid inclusions incorporated in the interlayer (Ellis et al., 2022a).

In most of the alkaline and calc-alkaline samples $K_{biotite}$ is followed by $K_{plagioclase}$ (see below), followed by $K_{clinopyroxene}$. In the alkaline trend clinopyroxene crystals from the hotspot setting (Pepom tephra) have the highest apparent partition coefficients (0.06–0.48) whereas those from the subduction-extensional setting have the lowest Kds (CI: 0.02–0.22, NYT: 0.02–0.18; Campi Flegrei: 0.02–0.22). For the calc-alkaline system the Kds of clinopyroxene crystals from subduction setting (Xoxoctic Tuff) vary between 0.09 and 0.40 while for orthopyroxene the range is from 0.15 to 0.17. The calculated apparent Kds for the alkaline pyroxene are in a good agreement with the values found by

<figure><img src="image_13.png" /><figcaption>**Fig. 9.** Lithium concentrations of the different mineral phases present in the alkaline trend and in the calc-alkaline trend as well as the Li concentrations in the co-existing glasses.</figcaption></figure>
<figure><img src="image_14.png" /><figcaption>**Fig. 10.** Partitioning of Li between co-existing minerals and glass organised by their tectonic settings and their chemical trend (alkaline vs. calc-alkaline). $Kd_{Li(min/melt)}$ are displayed for the different mineral phases. The left-hand side of the diagram shows the apparent partition coefficients ($Kd_{Li(min/melt)}$) for minerals (sanidine, plagioclase, biotite, clinopyroxene, quartz) co-existing with melts of phonolitic to trachyandesitic composition. On the right-hand side of the diagram apparent Kds for minerals co-existing with melts of rhyolitic to andesitic compositions are presented. Grey bar represents experimental data from Iveson et al. (2018).</figcaption></figure>

Beard et al. (2019). They reported Kd values from 0.16 to 0.43 with the lower value representing natural clinopyroxene–glass pairs from Tenerife, Canary Islands (hotspot setting). The calculated apparent Kds for the calc-alkaline pyroxene are always lower compared to the apparent Kds found by Neukampf et al. (2019) for rhyolitic ignimbrites from a hotspot setting (clinopyroxene Kds: 0.26–0.44; orthopyroxene Kds: 0.20–0.44) and experimentally determined Kd based on the analysis of

rhyodacitic glass (clinopyroxene $Kd = 0.26 \pm 0.04$) by Iveson et al. (2018). Olivine crystals from a subduction zone sample (one of the Xoxoctic Tuff samples) return lower Kds (0.04–0.15) compared to the published values (0.31–0.58) by Neukampf et al. (2019). However, the $Kd_{olivine}$ found in this study is in a good agreement with values determined by Grant and Wood (2010). Grant and Wood (2010) used a more forsteritic composition in their experiments yielding $Kd_{olivine}$ values
between 0.1 and 0.23, typical of mafic compositions, whereas $Kd_{olivine}$ calculated by Neukampf et al. (2019) are for more fayalitic composition (Fa$_{87-89}$). Olivine crystals from the subduction zone sample are Fo$_{61-77}$ and they are thus in a better agreement with $Kd_{olivine}$ appropriate for basaltic systems. Olivine crystals from another subduction zone sample (Xoxoctic Tuff sample from layer A) have a $Kd_{olivine}$ of 0.44. However, Li concentrations in the groundmass glass of this particular sample were enriched (> 100 ppm). In fact, all phases present in this sample display elevated Li contents and, by inference, yield higher Kds compared to those reported for the other samples.

### 5.1.2. Apparent partition coefficients for felsic phases

In the alkaline and calc-alkaline trend plagioclase crystals return a higher $Kd_{plagioclase}$ compared to $Kd_{sanidine}$. In the calc-alkaline samples either quartz or plagioclase always yield the highest Kd for the felsic mineral phases whilst sanidine always has the lowest Kd values. The quartz crystals from extensional tectonic settings (Caetano Tuff and BT) from the calc-alkaline trends return the highest apparent partition coefficients (0.1–0.9) whereas quartz crystals from a subduction zone setting (KPT) have the lowest Kds (0.09–0.13). The $Kd_{qtz}$ is on average higher compared to $Kd_{plag}$. However, the $Kd_{qtz}$ for subduction and extensional settings (BT and KPT) are lower compared to the published $Kd_{qtz}$ for the Mesa Falls Tuff (avg. rim: 0.41, avg. core: 0.45), published by Neukampf et al. (2019). The average $Kd_{plag}$ varies between 0.06 and 0.19 for the alkaline samples and from 0.03 to 0.55 for the calc-alkaline samples, which might reflect the period of a growth of plagioclase crystals in both systems. These values are consistent with the published partition coefficients ($Kd = 0.15–0.7$; Ryan, 1989; Blundy et al., 1998; Aigner-Torres et al., 2007; $Kd = 0.23 \pm 0.02$; Iveson et al., 2018; mean $Kd = 0.28$, Neukampf et al., 2019). The apparent $Kd_{sanidine}$ in the alkaline samples from hotspot to subduction-extension settings vary from 0.02 to 0.04 (CI, NYT, Astroni, Granadilla, Pepom tephra) with no difference between the two different types of tectonic settings and it varies from 0.05 (Jarbidge) to 0.10 (Bishop Tuff) in the calc-alkaline samples.

Previous studies (Ryan, 1989; Brenan et al., 1998a, 1998b; Taura et al., 1998; McDade et al., 2003; Ottolini et al., 2009) have shown that in mafic systems Li preferentially partitions into olivine ($Kd = 0.20–0.56$), followed by plagioclase ($Kd = 0.15–0.7$; Ryan, 1989; Blundy et al., 1998; Aigner-Torres et al., 2007) and to a lesser extent clinopyroxene ($Kd = 0.11–0.59$; Ryan, 1989; Hart and Dunn, 1993; Blundy et al., 1998; Brenan et al., 1998a, 1998b; McDade et al., 2003; Ottolini et al., 2009) and orthopyroxene ($Kd = 0.17–0.26$; Brenan et al., 1998a, 1998b; McDade et al., 2003; Ottolini et al., 2009). Neukampf et al. (2019) stated that for rhyolitic systems the order in which Li partitions into the mineral phases appears to mirror basaltic systems and the greatest abundances of Li can be found in fayalitic olivine, followed by oligoclase, clinopyroxene and orthopyroxene. Such an order is not observed for the calc-alkaline trend investigated in this study. Here, Li preferentially partitions into biotite followed by plagioclase, quartz, orthopyroxene, clinopyroxene and sanidine. We note here that we only determined the Li contents for olivine, plagioclase, clinopyroxene and orthopyroxene in samples from a subduction zone (Xoxoctic Tuff) whereas Neukampf et al. (2019) calculated the Kd values for mineral phases and co-existing groundmass glass for a total of 15 bodies from the Yellowstone-Snake River Plain and Mesa Falls Tuff (continental hotspot). We reinforce that the subduction zone samples (from Xoxoctic Tuff) are andesitic to dacitic in composition and not rhyolitic, thus representing an exception to the existing database; this may also influence the apparent distribution coefficients for the mineral phases. We find that, for the alkaline trend, when biotite is present Li preferentially partitions into biotite, followed by plagioclase, clinopyroxene and sanidine. The Kds may further be influenced by magmatic evolution as it exerts a dominant control on the major element composition of the crystallising mineral phases. Indeed, Beard et al. (2019) showed that the major element composition of clinopyroxene strongly influences the partition coefficients of elements. Additionally, the apparent partition coefficients may also be influenced

by the crystallisation sequence of the minerals and the concentrations of volatile species within the systems which means that, indirectly, the tectonic settings may play a significant role. However, this would need to be further investigated by means of experimental petrology.

## 5.2. Melt inclusions

For elements such as Li that are potentially mobilised into the vapour phase or brine under shallow magmatic conditions, melt inclusions are an important source of information about Li abundances of the melt during magmatic crystallisation because they can potentially entrap high Li contents in silicate melt prior to the processes that can strip Li from the melt phase. Despite a large number of melt inclusion studies that have been carried out in the past decades, there are relatively few that include Li, and for alkaline systems in particular, which is likely due to the lack of reliable host minerals in such systems. We have compiled available naturally glassy melt inclusion data, primarily hosted within quartz crystals, reflecting hotspot, subduction, and continental rift tectonic settings (Fig. 11). The subduction dataset includes results from settings where crustal influence is inferred to be significant (the Andes; e.g., Grock et al., 2017) as well as subordinate (Taupo; e.g., Liu et al., 2006). The hotspot and rift datasets are dominated by the Yellowstone and Valles systems, respectively. We exclude fully crystallised inclusions for reasons that are discussed below. The overarching conclusion from these data is that Li contents in melt inclusions are usually not orders of magnitude higher than those found in the co-existing groundmass, with partially degassed melt embayments commonly overlapping both the inclusion and groundmass fields.

While there are reports in the literature of erupted silicate melts containing high (>1000 ppm) Li contents (e.g., Taylor Creek: Webster and Duffield, 1991; Macusani: Pichavant and Montel, 1988; Hideaway Park, Spore Mountain: Benson et al., 2017) these are relatively rare. In the cases of Hideaway Park and Spore Mountain, these are petrographically unusual small-volume deposits (< 50 km³) and in the case of the Hideaway Park inclusions analysed by Mercer et al. (2015) were either crystallised or homogenised prior to analysis (discussed below). The Macusani obsidian, that is particularly Li-rich, is only found as lithic clasts within/weathered out from other volcanic deposits (Pichavant and Montel, 1988). The Taylor Creek rhyolite inclusions of Webster and Duffield (1991) are more intriguing as they represent a relatively standard subduction-type system from the Mogollon-Datil volcanic field (McIntosh et al., 1992). While still naturally glassy, we speculate that these inclusions may have experienced a slightly different post-eruptive cooling history within lava than the inclusions in rapidly quenched pyroclastic deposits illustrated in Fig. 10.

It has been widely recognised that melt inclusions remain the best way of quantifying pre-eruptive volatile contents of magmatic systems (Lowenstern, 1995). However, generating meaningful data from inclusions requires consideration of a range of processes that can modify inclusion composition. These can include those that are geological in origin such as post-entrapment crystallisation of the inclusion, fracturing of the inclusion (e.g., Wallace, 2005), or diffusion of components through (or indeed exchange with) the host mineral (Portnyagin et al., 2008; Lloyd et al., 2013). Superimposed on these issues are those that are imparted by laboratory processes. If melt inclusions are not rapidly quenched upon eruption, there exists the possibility for crystals to grow within the inclusion and if the cooling rate is sufficiently slow (as is the case in some portions of welded ignimbrites and the volumetric majority of lavas) the entire inclusion may be crystallised. Because many analytical techniques such as FTIR and SIMS require a homogeneous matrix, inclusions typically are re-homogenised under conditions thought to reflect their entrapment. However, such re-homogenisation is not without problems. In the best-case scenario, it may perfectly transform the crystalline inclusion to a glassy version of itself, yet if leakage of volatiles was the reason for the crystallisation of the inclusion, then even the re-homogenised inclusions may not yield meaningful
<figure><img src="image_15.png" /><figcaption>**Fig. 11.** Plot of Li vs. Rb concentrations in melt inclusions and re-entrants (melt embayments) from naturally glassy samples from explosive deposits from hotspot (Myers et al., 2016; Neukampf et al., 2019, 2022), subduction zone (Liu et al., 2006; Wright et al., 2012; Grock et al., 2017) and rift (Dunbar and Hervig, 1992; Wallace et al., 1999; Myers et al., 2019; Saalfield et al., 2022) settings. Rubidium is used as it reflects the melting and/or crystallisation of feldspars from melts.</figcaption></figure>

information. Alternatively, homogenisation of the inclusions at high temperatures and pressures may induce diffusion of mobile components as has been noted in a number of studies (e.g., Portnyagin et al., 2008; Hartley et al., 2015), and consequentially not provide magmatic compositions.

These potential modifications appear to be affecting the Li contents of melt inclusions severely. In Fig. 12, inclusions from the same deposit are compared with one set of glassy and one set of initially crystallised inclusions, either re-homogenised (Huckleberry Ridge Tuff) or analysed as crystallised (Mesa Falls Tuff). It is immediately apparent that the quenched and naturally glassy inclusions preserve lower Li contents than those that crystallised, yet within a single deposit, less mobile elements (in Fig. 12b,c Sr and Zr are used) show a concentration overlap between different inclusion types. Currently, there are only a handful of datasets that are available to make such comparisons possible and we highlight the importance of understanding this behaviour when estimating Li budgets of magmas.

### 5.3. Lithium isotope fractionation in evolved alkaline systems

Most stable isotopic systems, including Li, are dominated by equilibrium fractionation at high temperatures (Seitz et al., 2004; Magna et al., 2006b; Jeffcoate et al., 2007; Magna et al., 2008; Ionov and Seitz, 2008). The rapid diffusivity in many geologically relevant media, coupled with the large mass difference between ⁶Li and ⁷Li, renders Li particularly susceptible to kinetic isotopic fractionation which can ultimately erase the original equilibrium fractionation (e.g., Lundstrom et al., 2005; Parkinson et al., 2007). As yet, there are no experimentally derived equilibrium values for Li isotope compositions in minerals at magmatic conditions. From co-genetic suites of natural samples, some studies have suggested that magma evolution via fractional crystallisation causes a limited change in the ⁵⁷Li of the evolving magma over a range of MgO and SiO₂ compositions (Tomascak et al., 1999; Schuessler et al., 2009) although AFC processes for evolved alkaline systems may result in modification of ⁵⁷Li towards negative values (Hamelin et al., 2009; Ackerman et al., 2015). Here, using natural samples we assess how the differing cooling rates (immediately quenched, vitrophyric, slowly cooled and microcrystalline) can affect the Li inventory of a sample (Fig. 8) and the scales over which such a variability may be

observed. Any effect of alteration on the Li isotopic composition is not discussed in this paper, and only processes and samples that are unaltered are evaluated in the following.

The rapidly quenched MFT suite (continental hotspot setting) highlights that strong kinetic fractionation can be observed resulting in large intra-crystal ⁵⁷Li variations (e.g., plagioclase; Neukampf et al., 2020) as a consequence of syn-eruptive degassing. Simultaneously, the kinetic fractionation can lead to >10 ‰ variations in ⁵⁷Li for the groundmass glass with the ⁷Li-enriched composition representing more degassed portions of the groundmass glass (Neukampf et al., 2022), following theoretical predictions (Watson, 2017). When volcanic deposits are rapidly quenched, these large variations in isotopic compositions are preserved and post-eruptive mobility of Li can be largely excluded. The late-stage degassing and, consequently, the late-stage kinetic isotope fractionation overprints the magmatic ⁵⁷Li signal to a large extent and only the mineral cores might preserve the original magmatic signature. To infer the Li isotopic compositions of mantle sources, melt inclusions could be investigated although available data shows a significant spread of several tens per mil units (Gurenko et al., 2005; Schiavi et al., 2012) which could be attributed to dehydration occurring in the system. Nonetheless, the Li elemental and isotope inventory of melt inclusions can be overprinted due to degassing (Neukampf et al., 2022).

As stated above, equilibrium Li isotopic fractionation among mineral phases in evolved systems studied here remains under-explored. By using bulk mineral separates measured using conventional analytical techniques (MC-ICPMS) the information of intra-crystal variability in ⁵⁷Li may become obscured even though first-order information about the Li isotopic composition of the entire system under investigation can be obtained (e.g., Ackerman et al., 2015). We, therefore, recommend to analyse crystals in-situ, for example by Secondary Ionization Mass Spectrometry (SIMS), fs-LA-MC-ICPMS or micro-drilling coupled with MC-ICPMS measurements for obtaining Li isotopic compositions which are less modified by kinetic processes, in tandem with Li elemental analysis by means of LA-ICPMS. This combined in-situ approach may result in significantly more focused understanding of the Li behaviour at magmatic temperatures and how kinetic processes can influence the isotopic composition of different minerals. This in turn will allow to put better constrains on the processes that take place in the magma chamber.
<figure><img src="image_16.png" /><figcaption>**Fig. 12.** Plot of glassy and crystallised melt inclusions of Huckleberry Ridge Tuff Member A and Mesa Falls Tuff. (A) Plot of Li vs Rb for the two types of melt inclusions from Huckleberry Ridge Tuff (Myers et al., 2016; Benson et al., 2017) with crystallised melt inclusions being enriched in Li relative to the glassy melt inclusions. Here the glassy melt inclusions are re-homogenised. (B) Plot of Zr vs. Sr for the same melt inclusions shown in A. For Zr vs. Sr the melt inclusions have only small variability between the glassy and crystallised types. (C) Plot of Li vs Rb for crystallised and glassy melt inclusions from Mesa Falls Tuff (Neukampf et al., 2019, 2022). The Li concentrations of the glassy melt inclusions are dramatically lower compared to the re-homogenised glassy melt inclusions in A. (D) Plot of Zr vs. Sr for both types of melt inclusions showing no variations between the two types.</figcaption></figure>

Another potential source of kinetic fractionation, that can be observed in rapidly quenched deposits, is the pre-eruptive late-stage fluxing of an exsolved Li-rich/Li-poor vapour phase as observed for Mount St. Helens (USA; Berlo et al., 2004; Kent et al., 2007). Berlo et al. (2004) and Kent et al. (2007) argued that the Li enrichments in plagioclase by a factor of two compared to material erupted subsequently from pumice to dome samples must be caused by exsolved alkali-enriched vapour in the upper part of the magma chamber. They, additionally, investigated plagioclase crystals from ash samples that have a lower Li concentration which they interpreted as a result of degassing. We note that the Li elemental variation observed in their samples could equally be a result of post-eruptive Li mobility whereby slow cooling of the dome lava may lead to an enrichment in Li concentration due to groundmass crystallisation (discussed further below). However, fluxing of a Li-rich vapour phase would lead to Li enrichment in the rims of the crystals and to a lighter isotopic composition in the rims because ⁶Li diffuses faster into the crystals relative to ⁷Li (Parkinson et al., 2007; Richter et al., 2009; Holycross et al., 2018). A Li-poor vapour phase would lead to the opposite effect of lower Li concentration in the rims, paralleled by a heavy isotopic composition, which is

similar to what can be observed for late-stage degassing. Magma mixing and crystal fractionation can result in differences in Li concentration and isotopic composition in minerals, but these changes would be reflected by other major and trace elements.

The δ⁷Li composition for the bulk quenched samples from this study (MFT, KPT, BT, Astroni, Granadilla) varies by ~10 ‰. This difference in Li isotopic composition is unrelated to the melt composition (rhyolitic vs. phonolitic). The bulk δ⁷Li values always represent an accumulated history of the sample, i.e. from the mantle source through potential assimilation-crystal fractionation to degassing. Therefore, using bulk δ⁷Li isotopic data will only be meaningful when all processes that can influence the isotopic systematics are considered, addressed and quantified. Of the samples considered here only MFT bulk-rock δ⁷Li values (3.9–4.3 ‰; Neukampf et al., 2019) are within the range of published MORB data (4 ± 2 ‰; Elliott et al., 2006; Tomascak et al., 2008; Marshall et al., 2017) whereas δ⁷Li values in all other samples strongly deviate from the published MORB range. This highlights the fact that bulk rock δ⁷Li values of samples that either underwent a protracted period of cooling or significant degassing should not be compared to MORB as they do not represent the equilibrium δ⁷Li values of these
systems. Marschall et al. (2017) observed that MORB glasses that have $\delta^7$Li values higher than +4.0 ‰ reflect assimilation of low-T altered oceanic crust into the magma chamber. They also observed $\delta^7$Li variations of 1.0–1.5 ‰ among samples from different mid-ocean ridge segments and attributed this 1.0–1.5 ‰ variation to kinetic fractionation of Li isotopes during magma extraction, transport and storage.

For clinopyroxene separates there is no large $\delta^7$Li variation between the rhyolitic (−0.1‰ to +1.0 ‰, MFT; Neukampf et al., 2019) and phonolitic (0.5 ‰; Astroni) samples. A similar observation can be made for plagioclase when considering the data solely obtained using MC-ICPMS. For rhyolitic and phonolitic samples the differing Li concentrations in plagioclase (13 ppm vs. 3 ppm, respectively) and clinopyroxene (11 and 14 ppm vs. 8 ppm) are not reflected in their Li isotopic compositions. At least for clinopyroxene it could be argued that, independent of concentration and cooling environment (rapidly quenched vs. vitrophyre), the $\delta^7$Li values are identical within ±1.5 ‰.

For vitrophyric samples that remained sufficiently hot for post-emplacement sintering and welding, but cooled fast enough to produce a glassy groundmass, Ellis et al. (2018) reported a limited $\delta^7$Li range (+5.2 to +7.6 ‰) for the Tuff of Knob rhyolite (continental hotspot). The $\delta^7$Li values for the mineral separates of the dense vitrophyre from the Campanian Ignimbrite sample (Pro 4–1) from the subduction-extensional setting show a similar trend and have a $\delta^7$Li range of 10 ‰ (from biotite to groundmass glass), but we note here that the mineralogy of the samples is very different to those of the Tuff of Knob. The Caetano Tuff (extensional setting) spans a $\delta^7$Li range of 21 ‰ (from biotite to quartz). Similar to the quenched deposits the vitrophyre should retain some of the magmatic signal, but may also experience kinetic fractionation during eruption and/or due to cooling after the eruption although no in-situ data is yet available. When considering the three examples representative of a quickly cooled vitrophyre in Fig. 8 it appears that the phonolitic sample (CI) is isotopically lighter compared to the rhyolitic counterpart. However, this paucity of data from evolved alkaline volcanic rocks and their constituent minerals only serves to highlight that, compared to more traditional isotopic systems (e.g., Sr, Nd, Pb, O), addressing major processes taking place post-eruption using Li isotopic systematics is still premature.

Ellis et al. (2018) showed that for a continental hotspot the Li concentrations in the mineral phases (plagioclase, sanidine, augite, pigeonite) from eight different deposits are notably elevated due to groundmass crystallisation during the cooling of the ignimbrite. This is verified by comparing the vitrophyric Caetano Tuff sample to the devitrified Caetano Tuff sample from this study. During cooling Li diffuses into the crystals and results in Li enrichment. As a consequence of this post-eruptive mobility the original magmatic signal of Li in the minerals will mostly be overprinted. Using fs-LA-MC-ICPMS, Steinmann et al. (2020) showed that olivine crystals from the Roche Sauterre lava flow/lava lake, the Massif Central (France), retained their intrinsic magmatic Li signal only in their cores, but olivine rims have shown a uniform increase in Li concentration, paralleled by isotopically heavier Li.

The Li isotope systematics in quartz appear to be independent of the cooling environment. The quartz separates yield the isotopically heaviest values of the mineral separates, otherwise the groundmass glass carries the highest values. In quartz Li is tetrahedrally coordinated (Soltay and Henderson, 2005) and, therefore, quartz favours $^7$Li over $^6$Li making it isotopically heavier than other common mineral phases (Teng et al., 2006; Magna et al., 2016). Quartz crystals from the Bishop Tuff (10.1‰ and ~ 16.7 ppm), Caetano Tuff (10.0–11.5‰ and 11.4–20.9 ppm) and Mesa Falls Tuff (6.7–7.2‰ and 17.1–18.8 ppm) have similar Li abundances and comparable isotopic compositions whereas quartz crystals from the Kos Plateau Tuff have significantly lower Li concentrations (mean: 8.9 ppm) and are isotopically lighter ($\delta^7$Li ~ 2.0‰) compared to the other eruption units. This difference in abundances and isotopic compositions could be the result of the difference in water concentration ($OH^-$ and $H_2O$) in the crystal structure for the different

eruption units. Since $Li^+$ and $H^+$ charge balance $Al^{3+}$ ($Si^{4+}$ ↔ $Al^{3+}$ + $Li^+$ or $H^+$) in the quartz crystal structure (e.g., Jollands et al., 2020), $Li^+$ and $H^+$ are competing for the same site in quartz (e.g., Tollan et al., 2019). Therefore, at higher $H^+$ concentration less $Li^+$ can be incorporated into the crystal structure. From Fig. 3 it is apparent that KPT (subduction zone samples) and BT (extensional setting sample) plot in the Al-excess region suggesting that $Li^+$ is charge balancing Al insufficiently and that $H^+$ is required to reduce the charge imbalance. In contrast, the Caetano Tuff samples (extensional setting) plot on and around the Al:Li equity line indicating a sufficient charge balance of Al by $Li^+$ and no $H^+$ is thus needed. However, with the data available up to date it is not possible to provide unequivocal proof for this theory. Jollands et al. (2020) argue that when $Li^+$ and $H^+$ are exchanging and the elapsed time is not sufficiently long to eliminate the H-Li gradient entirely, a diffusion profile should be observed. This means that there could be kinetic Li isotopic fractionation due to diffusion in quartz which would form an Li isotopic gradient within quartz. However, this demands additional testing using in-situ $\delta^7$Li measurements. The Caetano Tuff quartz samples prove that Li concentration in quartz is not significantly affected by the type of sample (glassy vs. devitrified) which has previously been shown by Ellis et al. (2018).

## 6. Conclusions

The extensive compilation of data for different magmatic systems allows us to define the Li budget for these systems and consider the potential magmatic and post-magmatic controls on Li distribution in these systems. We provide evidence that Li is incompatible in all mineral phases crystallising during magmatic differentiation that typifies most global volcanism. The Li concentration for glass and co-existing minerals for the alkaline trend decreases as following: groundmass glass (1.1–318 ppm), biotite (0.8–99.3 ppm), plagioclase (0.9–25.1 ppm), clinopyroxene (0.6–31.0 ppm), sanidine (0.5–38.7 ppm). For the calc-alkaline trend Li decreases as following: biotite (10.2–2314 ppm), groundmass glass (2.1–176 ppm), plagioclase (1.4–35.0 ppm), quartz (3.2–21.1 ppm), olivine (1.0–16.3 ppm), clinopyroxene (0.7–21.8 ppm), sanidine (0.3–8.1 ppm). Although the highest Li concentrations can be found in the groundmass glass, these concentrations are not high enough to form economically recoverable deposits (e.g., Benson et al., 2017).

Our data shows that for evolved magmas the Li concentration in glass and minerals is independent of the tectonic setting, but more closely depends on cooling environment. Therefore, minerals from slowly cooled ignimbrites and lavas from evolved melts are enriched in Li relative to quickly quenched/cooled portions of the same deposits. However, this still needs to be tested systematically for mafic systems. We found that for the alkaline trend when biotite is present Li preferentially partitions into biotite followed by plagioclase, clinopyroxene and sanidine. In the calc-alkaline samples Li preferentially partitions into biotite followed by plagioclase, quartz (potentially modulated by the water content of the melt), orthopyroxene, clinopyroxene and sanidine.

The Li concentrations in melt inclusions appear to be independent of tectonic setting. However, the re-homogenisation of crystallised melt inclusions may result in elevated Li contents in the inclusions for reasons that as yet remain unclear. This observation requires further studies as this will have a major effect on estimating the Li budget in magmatic reservoirs and processes involving concentrating Li to economically important levels.

Crystals from rapidly quenched deposits may show large intra-crystal $\delta^7$Li variations due to syn-eruptive processes (degassing or vapour release from deeper parts of the magma chamber) that lead to kinetic fractionation and only mineral cores have the potential to preserve any original magmatic signal. Likewise, the groundmass glass will likely record a large variability in $\delta^7$Li. Slowly cooled deposits have a large range in $\delta^7$Li values (> 10 ‰) and rapidly cooled vitrophyre deposits have a similar range in Li isotopic composition. The $\delta^7$Li variations of
rapidly quenched to slowly cooled deposits most likely represent disequilibrium/kinetic fractionation and only the cores of minerals in all three tectonic settings might represent the intrinsic isotopic signature of the magma chamber. We observed no difference in Li isotopic composition between phonolitic versus rhyolitic samples. Thus, the only differences are observed for the Li concentrations of the different minerals. The intra-crystal record of Li isotopes in such chemically evolved alkaline silicate systems therefore remains largely unexplored and such studies have significant future potential. Therefore, further investigations with in-situ methods are needed to evaluate the Li isotopic behaviour in single crystals from slowly cooled ignimbrites and lavas as well as quickly cooled vitrophyres.

## Declaration of Competing Interest

The authors declare the following financial interests/personal relationships which may be considered as potential competing interests:
Dr. Ben Ellis reports financial support was provided by Swiss National Science Foundation.

## Data availability

All data are available in sup. materials.

## Acknowledgements

This work was supported by the Swiss National Science Foundation (grant 200021_166281 to BE). TM acknowledges support from the Czech Science Foundation project 19-29124. The manuscript benefited from detailed reviews by two anonymous reviewers. Claudia Romano is acknowledged for editorial handling.

## Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi.org/10.1016/j.chemgeo.2023.121628.

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