# A compositional tipping point governing the mobilization and eruption style of rhyolitic magma

D. Di Genova, S. Kolzenburg, S. Wiesmaier, E. Dallanave, D. R. Neuville, K.

U. Hess, D. B. Dingwell

► To cite this version:

D. Di Genova, S. Kolzenburg, S. Wiesmaier, E. Dallanave, D. R. Neuville, et al. A compositional tipping point governing the mobilization and eruption style of rhyolitic magma. Nature, 2017, 552, pp.235-238. 10.1038/nature24488. insu-03748823

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# A compositional tipping point governing the mobilization and eruption style of rhyolitic magma

D. Di Genova<fnref n="1" />,², S. Kolzenburg², S. Wiesmaier², E. Dallanave², D. R. Neuville³, K. U. Hess² & D. B. Dingwell²

The most viscous volcanic melts and the largest explosive eruptions<fnref n="1" /> on our planet consist of calcalkaline rhyolites²,³. These eruptions have the potential to influence global climate⁴. The eruptive products are commonly very crystal-poor and highly degassed, yet the magma is mostly stored as crystal mushes containing small amounts of interstitial melt with elevated water content⁵. It is unclear how magma mushes are mobilized to create large batches of eruptible crystal-free magma. Further, rhyolitic eruptions⁶⁻⁸ can switch repeatedly between effusive and explosive eruption styles and this transition is difficult to attribute to the rheological effects of water content or crystallinity⁹,¹⁰. Here we measure the viscosity of a series of melts spanning the compositional range of the Yellowstone volcanic system and find that in a narrow compositional zone, melt viscosity increases by up to two orders of magnitude. These viscosity variations are not predicted by current viscosity models¹¹,¹² and result from melt structure reorganization, as confirmed by Raman spectroscopy. We identify a critical compositional tipping point, independently documented in the global geochemical record of rhyolites, at which rhyolitic melts fluidize or stiffen and that clearly separates effusive from explosive deposits worldwide. This correlation between melt structure, viscosity and eruptive behaviour holds despite the variable water content and other parameters, such as temperature, that are inherent in natural eruptions. Thermodynamic modelling demonstrates how the observed subtle compositional changes that result in fluidization or stiffening of the melt can be induced by crystal growth from the melt or variation in oxygen fugacity. However, the rheological effects of water and crystal content alone cannot explain the correlation between composition and eruptive style. We conclude that the composition of calcalkaline rhyolites is decisive in determining the mobilization and eruption dynamics of Earth's largest volcanic systems, resulting in a better understanding of how the melt structure controls volcanic processes.

The structure and viscosity of natural rhyolites are known to be greatly affected by the ratio of alkalis and alkaline earths to aluminium and iron¹³. Raman spectra of synthetic rhyolites¹⁴ show rearrangements of the melt structure caused by Na-K substitution, resulting in a substantial viscosity change. However, no systematic studies of these effects in natural melts have been presented so far.

Hence, we performed an experimental investigation of the effect of small compositional changes on the viscosity of rhyolites. These compositional changes may result from magma crystallization, mixing or reheating. We carried out viscosity measurements on 11 rhyolite samples, mimicking the compositional range of the Yellowstone volcanic system, which exhibits both effusive and explosive eruptive behaviour. We describe the chemistry of the samples (Extended Data Table 1) using a new melt structure index, termed the rheological agpaitic index (RAI; calculated in mol%), and defined as:

$$ \text{RAI} = (\text{Na}_2\text{O} + \text{K}_2\text{O} + \text{CaO} + \text{MgO} + \text{FeO})/(\text{Al}_2\text{O}_3 + \text{Fe}_2\text{O}_3) \quad (1) $$

RAI characterizes the transition of the structural role of alkalis, alkaline earth elements and ferrous iron from charge balancing (RAI < 1) to network modifying (RAI > 1) (see Methods). Additional parameters for the characterization of the samples are the molar ratio of K₂O/(K₂O + Na₂O) (hereafter termed K#) and the amount of iron dissolved in the melt.

RAI indices in the first nine samples range from 0.75 (charge balanced) to 1.12 (network modified) and K# ranges from 0.19 to 0.77 (samples A-I). To investigate the effect of magma mixing, a well documented process at Yellowstone¹⁵ that mainly results in iron enrichment and relative depletion of silica, two additional samples (with K# ≈ 0.5, samples J and L) were synthesised. These samples have RAI values of up to 1.54, namely, a more depolymerized structure. The chemical compositions of all samples, along with K# and RAI, are reported in Extended Data Table 1.

Melt viscometry between 750 and 1,590 °C isolated the effects of RAI, K# and iron/silica content on viscosity (Extended Data Fig. 1 and Extended Data Table 2). Increasing RAI (at K# ≈ 0.5 and constant temperature of 850 °C) results in a decrease in melt viscosity by up to about 1.5 log units (Fig. 1a and b). This change occurs abruptly at RAI = 1 (Fig. 1b) and results from structural reorganization due to the complete charge balancing of network formers (RAI ≤ 1), forcing remnant cations into a network modifier role (RAI > 1, see Methods). Notably, the shift in RAI across a value of 1 is caused by a decrease of 0.5 wt% of Al₂O₃ (for example, samples C and H, with RAI values 0.92 and 1.07, respectively). For samples free of network modifiers (RAI < 1), changes in K# correlate linearly with viscosity and result in variations of up to 1.5 logarithmic units (Fig. 1a). In such a structural configuration, substitution of Na by K increases rhyolite viscosity and vice versa. Importantly, current models of magma viscosity are unable to identify both RAI and K# effects on viscosity (see Methods and Extended Data Fig. 2).

To investigate the effect of magma mixing, a process acting on a timescale as short as minutes to hours¹⁵, the two FeO-rich SiO₂ depleted samples (J and L in Extended Data Table 1) are compared with an iron-poor sample (F). At 850 °C, the iron-poor SiO₂-rich sample (FeO = 1.55 wt%) is the most viscous (Fig. 2). Increasing iron content from 1.55 wt% to 3.31 wt%, together with decreasing silica by about 5 wt%, results in a substantial reduction of viscosity (about 2.5 orders of magnitude) through structural depolymerization. However, further increasing FeOtot to 6.24 wt% and decreasing SiO₂ by 2.5 wt% does not result in a further viscosity decrease.

Instead, the iron-rich sample (L) shows an extremely high viscosity, similar to the iron-poor sample (F), in contrast to the expected¹¹,¹² decrease in viscosity, and results in a difference of four orders of magnitude between current models and the measured value. We relate this viscosity increase to the growth of FeO nanocrystals due to iron oversaturation in the melt. This effectively depletes the melt structure from iron, resulting in structural polymerization (see Methods).

Raman spectra and magnetic-hysteresis analyses of sample L identified these crystals as magnetite with diameters between about 5 nm and 30 nm (Methods and Extended Data Figs 3 and 4). Differential

<footnote n="1">¹School of Earth Sciences, University of Bristol, Wills Memorial Building, Queens Road, Bristol BS8 1RJ, UK. ²Department of Earth and Environmental Sciences, Ludwig-Maximilians-Universität München, Theresienstraße 41/III, 80333 München, Germany. ³Géomatériaux, CNRS-IPGP UMR 7154, Paris Sorbonne Cité, Paris, France.</footnote>
<figure><img src="image_2.png" /><figcaption>**Figure 1** | Viscosity measurements of samples at 850 °C. a, Measured viscosity for samples A to I characterized by similar SiO₂ content (77.32 ± 1.03 wt%; see Extended Data Tables 1 and 2 for chemical composition and viscosity measurements, respectively) as a function of K# (calculated as K₂O/(K₂O + Na₂O)). Labels indicate RAI for each sample (Extended Data Table 1). The viscosity increases with increasing K₂O content. b, Measured viscosity as a function of RAI for samples characterized by similar SiO₂ content (77.03 ± 0.46 wt%) and K#.

Labels indicate K# for each sample. The viscosity abruptly decreases at RAI > 1. The error in the viscosity measurements was estimated to be ±0.05 logarithmic units (see Methods).</figcaption></figure>

thermal analysis further corroborated this effect. The iron-rich sample (L) shows two endothermic peaks (Extended Data Fig. 5), reflecting the presence of two domains in the glass structure, nanolite-free and nanolite-bearing. The differential thermal analysis data allow us to calculate the viscosity of the nanolite-free domain (L-DTA in Fig. 2 and Methods). The viscosity of the nanolite-free melt (that is, fully dissolved, high FeO content) is the lowest of all measured viscosities (Fig. 2).

In summary, the viscometry shows that rhyolite viscosity decreases with increasing RAI, decreasing K# and when FeO is substituted for SiO₂. This substitution effect is in agreement with existing models and literature data¹¹,¹². However, this only holds up to the saturation point of iron, beyond which the precipitation of nanolite crystals dramatically increases melt viscosity.

Pre- and syn-eruptive variation of RAI, K# and the FeO content and its oxidation state can occur through open- or closed-system processes and on varying timescales¹⁰,¹⁵⁻²³. Open-system processes, such as recharge or reheating of a magma reservoir or crystal mush, magma mixing, crustal assimilation, and conduit processes, may greatly change the magma composition and redox conditions of the system. To constrain the impact of closed-system processes, we simulated both equilibrium and fractional crystallization along the liquid line of descent of a hydrous (H₂O = 1 wt%) rhyolitic melt with K# of 0.5 (that is, containing the same molar amount of Na₂O and K₂O) at 0.5 kbar and 1.5 kbar (Methods, Supplementary Table 1) from the liquidus conditions down

<figure><img src="image_3.png" /><figcaption>**Figure 2** | Measured viscosity at 850 °C for samples F, J and L characterized by increasing FeO content. FeO content (in wt%) is given in parentheses; see Extended Data Table 1. Arrows show the increase in measured iron content starting from sample F. However, an oversaturation of the system results in the formation of Fe-oxide nanolites (sample L), leading to an effective FeO depletion of the melt and a viscosity hysteresis back to higher viscosities. Using differential thermal analysis data (Methods) we calculated the nanolite-free viscosity (L-DTA) of sample L. The error in the viscosity measurements was estimated to be ±0.05 logarithmic units (see Methods).</figcaption></figure>

to 750 °C using Rhyolite-MELTS²⁴. The simulations were performed at oxygen fugacities ($f_{O_2}$) representative of magmatic conditions (QFM and QFM + 2 buffer; Methods). With increasing degree of crystallization, the model predicts decreasing RAI and increasing K# (Fig. 3a and b), resulting in increasing viscosity. At QFM + 2, less than 10% crystallization is required to cross the rheological tipping point of RAI ≈ 1 (Fig. 3a). Further, RAI increases with decreasing oxygen fugacity $f_{O_2}$ (Fig. 3a, from QFM + 2 to the QFM buffer condition), because of the bivalent structural role of Fe in the melt: both as network-former Fe³⁺ and network-modifier Fe²⁺; see equation (1). K# increases with decreasing pressure (Fig. 3b, from 1.5 kbar to 0.5 kbar). Shallower magma bodies may thus reach higher K# and consequently higher viscosities. The modelling indicates the capacity of closed-system processes to change magma viscosity on relatively short timescales.

These results show that changes in RAI and K# correlate with drastic variations in viscosity and may, therefore, influence the eruptive behaviour of rhyolitic melts on short timescales. To test this observation, we plot RAI versus K# for all rhyolitic systems where matrix glass compositions are available (Yellowstone and 40 other volcanic systems, totalling more than 950 chemical analyses). We find that RAI and K# systematically correlate with the eruptive styles of the respective deposits (see also Methods and Supplementary Table 2). Effusive and explosive products, indistinguishable in a total alkali versus silica diagram (Extended Data Fig. 6), separate into two distinct groups (Fig. 4) at RAI > 1 (effusive) and RAI < 1 (explosive), defined by the experimentally determined rheological tipping point at RAI = 1.

Detailed analysis of the Yellowstone data shows that explosive products are characterized by K# > 0.5, and effusive samples by K# < 0.5, indicating also a contribution of coupling of the RAI and K# effects. A similar trend is observed in the data for explosive and effusive phases of the same volcano (open symbols in Fig. 4). Overall, the experimentally identified tipping point is reflected in the geochemical record of rhyolitic systems across the globe.

Previous explanations for explosive–effusive transitions, based largely on variations in temperature, crystallinity and volatile content of the magma, have failed to fully explain the separation of explosive and effusive eruptive behaviour. We acknowledge that these parameters have a fundamental role in determining the intensity of an eruption through their effect on magma viscosity and because volatiles supply the main driving force for eruptions. However, through careful isolation
<figure><img src="image_4.png" /><figcaption>**Figure 3** | Thermodynamic modelling results of magma crystallization at varying $f_{O_2}$ and pressure. a, Modelled RAI using sample composition D (Extended Data Table 1). b, Modelled K# using sample composition D. Simulations were performed using the Rhyolite-MELTS²⁴ thermodynamic software. Outputs at 0.5 kbar and 1.5 kbar are simulated with $f_{O_2}$ buffered at QFM and QFM + 2 (more oxidized) and equilibrium crystallization conditions. Crystallization strongly decreases RAI and increases K#.</figcaption></figure>

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

At fixed temperature, a more reduced condition (QFM) increases RAI (Fig. 3a), while pressure increases K# at low temperature (Fig. 3b). Black circles show the residual melt in wt% (for each 10%), while +qtz, +alb, +mag, +san, +opx and +degassing indicate the onset of quartz, albite, magnetite, sanidine and orthopyroxene crystallization and melt degassing, respectively. The chemical compositions of the residual melts and solid phases together with the phase assemblage are reported in Supplementary Table 1.

of melt structural effects at constant environmental parameters (volatile content and $f_{O_2}$), we find that the explosive-effusive divide is decisively controlled by small changes in the melt structure of rhyolites that are fundamentally affected by melt composition.

The third effect identified here (iron depletion in the melt caused by nanolite formation) may play an important part in the eruptive dynamics of not only rhyolitic but also lower-silica melts. Recent advances in micro-analytical techniques identified iron-bearing nanoparticles at

<figure><img src="image_6.png" /><figcaption>**Figure 4** | Summary plot of RAI versus K# for all experimental samples along with 40 natural rhyolitic systems and their eruptive styles.</figcaption></figure>

a, Density distribution map (hot colours indicate higher density; cold colours indicate lower density) of 236 compositional analyses of the glass matrix of lava flows and pyroclastic rocks of the Yellowstone volcanic system (Supplementary Table 2) and the experimental samples (black squares) plotted as RAI versus K#. Explosive and effusive products clearly separate into two distinct populations characterized by different RAI indices. These two fields depict two distinct melt rheology fields: a low-viscosity region (RAI > 1), where effusive deposits cluster, and a

high-viscosity region (RAI < 1), where explosive deposits cluster. b, The drastic change in eruptive behaviour at a well defined change point in melt structure (RAI = 1) is portrayed in the plot of RAI versus K# for a plethora of SiO₂-rich volcanic centres on Earth (40 volcanic systems; about 730 chemical analyses). Open symbols represent data from the same volcano. The inset shows the same data divided by eruption style only (explosive in red; effusive in blue). This confirms the universal nature of this change point in melt structure. Dashed lines represent RAI variation due to analytical uncertainty (see Methods).
the melt-bubble interface in the 2013 paroxysmal activity of Mount Etna (Italy) and linked them to unexpected increases in eruption intensity²⁵. This suggests that these particles may provide sites for the heterogeneous nucleation of bubbles. A recent study²⁶ distinguishes products from explosive-effusive transitions of the 2011 eruption of Shinmoedake (Japan) by the presence of nanolites. Thermal analysis²⁷ performed on samples derived from a violent Strombolian eruption at Mount Yasur Volcano (Vanuatu) suggested the presence of two melt domains as described and analysed in this study. Hence, this process appears not to be restricted to rhyolitic magmas; it also takes place in magmas with lower SiO₂ content.

In conclusion, experimental data, thermodynamic modelling and analysis of compositional data from the global volcanic record provide combined evidence for a sudden rheological transition in rhyolite. We observe three effects:

(1) Increasing RAI to values larger than 1 triggers a sharp decrease in viscosity by up to two orders of magnitude, representing a strict and sharp transition, separating effusive from explosive eruptions.

(2) Increasing K# results in a continuous, gradual increase in viscosity (at RAI < 1). However, explosive eruptions consistently display higher K# than their effusive counterparts.

(3) Melt viscosity decreases with increasing iron content and decreasing silica content until iron saturation is reached, beyond which iron depletion of the melt due to nanolite formation results in a drastic increase of viscosity.

The identified viscosity discontinuity can be crossed by small compositional changes in rhyolitic magmas and can be induced by crystallization, assimilation, magma replenishment or mixing. Composition-induced viscosity variations may also originate from changes in magma intrinsic parameters such as temperature, pressure or $f_{O_2}$.

The RAI threshold represents a composition-dependent viscosity tipping point in rhyolitic systems. This transition plays a key part in determining the erupitive style of a magma because a felsic magma or mush may be revitalized via chemical fluidification (RAI increasing to >1, low K# and iron enrichment) and because a viscosity increase (RAI < 1, high K# and iron depletion) may hinder efficient rhyolite degassing and lead to increased explosive potential. Further, independently of which process is acting during the generation of effusive rhyolitic deposits (that is, whether the material was never fragmented²⁸ or whether it underwent subsurface fragmentation and re-welding before extrusion²⁹,³⁰), RAI clearly separates the two eruptive styles on the basis of the melt structure.

**Online Content** Methods, along with any additional Extended Data display items and Source Data, are available in the online version of the paper; references unique to these sections appear only in the online paper.

**Received 27 March; accepted 27 September 2017.**

1. Bryan, S. E. et al. The largest volcanic eruptions on Earth. *Earth Sci. Rev.* **102**, 207–229 (2010).

2. Friedman, I., Long, W. & Smith, R. L. Viscosity and water content of rhyolite glass. *J. Geophys. Res.* **68**, 6523–6535 (1963).

3. Shaw, H. R. Viscosities of magmatic silicate liquids: an empirical method of prediction. *Am. J. Sci.* **272**, 870–893 (1972).

4. Robock, A. et al. Did the Toba volcanic eruption of ~74 ka B.P. produce widespread glaciation? *J. Geophys. Res. Atmos.* **114**, D10107 (2009).

5. Bachmann, O. & Bergantz, G. W. On the origin of crystal-poor rhyolites: extracted from batholithic crystal mushes. *J. Petrol.* **45**, 1565–1582 (2004).

6. Christiansen, R. L. The Quaternary and Pliocene Yellowstone plateau volcanic field of Wyoming, Idaho, and Montana. *U. S. Geol. Surv. Prof. Pap.* **729-G**, 145 (2001).

7. Castro, J. M., Manga, M. & Cashman, K. V. Dynamics of obsidian flows inferred from microstructures: insights from microlite preferred orientations. *Earth Planet. Sci. Lett.* **199**, 211–226 (2002).

8. Carn, S. A. et al. The unexpected awakening of Chaitén volcano, Chile. *Eos* **90**, 205–206 (2009).

9. Cabrera, A., Weinberg, R. F. & Wright, H. M. N. Magma fracturing and degassing associated with obsidian formation: the explosive-effusive transition. *J. Volcanol. Geotherm. Res.* **298**, 71–84 (2015).

10. Befus, K. S. & Gardner, J. E. Magma storage and evolution of the most recent effusive and explosive eruptions from Yellowstone Caldera. *Contrib. Mineral. Petrol.* **171**, 30 (2016).

11. Hui, H. J. & Zhang, Y. Toward a general viscosity equation for natural anhydrous and hydrous silicate melts. *Geochim. Cosmochim. Acta* **71**, 403–416 (2007).

12. Giordano, D., Russell, J. K. & Dingwell, D. B. Viscosity of magmatic liquids: a model. *Earth Planet. Sci. Lett.* **271**, 123–134 (2008).

13. Mysen, B. O. & Toplis, M. J. Structural behavior of Al³⁺ in peralkaline, metaluminous, and peraluminous silicate melts and glasses at ambient pressure. *Am. Mineral.* **92**, 933–946 (2007).

14. Le Losq, C. & Neuville, D. R. Effect of the Na/K mixing on the structure and the rheology of tectosilicate silica-rich melts. *Chem. Geol.* **346**, 57–71 (2013).

15. Morgavi, D. et al. The Grizzly Lake Complex (Yellowstone Volcano, USA): mixing between basalt and rhyolite unravelled by microanalysis and X-ray microtomography. *Lithos* **260**, 457–474 (2016).

16. Vazquez, J. A., Kyriazis, S. F., Reid, M. R., Sehler, R. C. & Ramos, F. C. Thermochemical evolution of young rhyolites at Yellowstone: evidence for a cooling but periodically replenished postcaldera magma reservoir. *J. Volcanol. Geotherm. Res.* **188**, 186–196 (2009).

17. Girard, G. & Stix, J. Rapid extraction of discrete magma batches from a large differentiating magma chamber: the Central Plateau Member rhyolites, Yellowstone caldera, Wyoming. *Contrib. Mineral. Petrol.* **160**, 441–465 (2010).

18. Wiesmaier, S. et al. Bimodality of lavas in the Teide-Pico Viejo succession in Tenerife—the role of crustal melting in the origin of recent phonolites. *J. Petrol.* **53**, 2465–2495 (2012).

19. Bachmann, O., Deering, C. D., Lipman, P. W. & Plummer, C. Building zoned ignimbrites by recycling silicic cumulates: insight from the 1,000km³ Carpenter Ridge tuff, CO. *Contrib. Mineral. Petrol.* **167**, 1025 (2014).

20. Ellis, B. S., Bachmann, O. & Wolff, J. A. Cumulate fragments in silicic ignimbrites: the case of the Snake River Plain. *Geology* **42**, 431–434 (2014).

21. Stelten, M. E., Cooper, K. M., Vazquez, J. A., Calvert, A. T. & Glessner, J. J. G. Mechanisms and timescales of generating eruptible rhyolitic magmas at Yellowstone caldera from zircon and sanidine geochronology and geochemistry. *J. Petrol.* **56**, 1607–1642 (2015).

22. Wolff, J. A. et al. Remelting of cumulates as a process for producing chemical zoning in silicic tuffs: a comparison of cool, wet and hot, dry rhyolitic magma systems. *Lithos* **236–237**, 275–286 (2015).

23. Gaillard, F., Scaillet, B. & Pichavant, M. Kinetics of iron oxidation-reduction in hydrous silicic melts. *Am. Mineral.* **87**, 829–837 (2002).

24. Gualda, G. A. R., Ghiorso, M. S., Lemons, R. V. & Carley, T. L. Rhyolite-MELTS: a modified calibration of MELTS optimized for silica-rich, fluid-bearing magmatic systems. *J. Petrol.* **53**, 875–890 (2012).

25. Barone, G. et al. Nanoscale surface modification of Mt. Etna volcanic ashes. *Geochim. Cosmochim. Acta* **174**, 70–84 (2016).

26. Mujin, M. & Nakamura, M. A nanolite record of eruption style transition. *Geology* **42**, 611–614 (2014).

27. Kremers, S. et al. Shallow magma-mingling-driven Strombolian eruptions at Mt Yasur volcano, Vanuatu. *Geophys. Res. Lett.* **39**, L21304 (2012).

28. Eichelberger, J. C., Carrigan, C. R., Westrich, H. R. & Price, R. H. Non-explosive silicic volcanism. *Nature* **323**, 598–602 (1986).

29. Gonnermann, H. M. & Manga, M. Explosive volcanism may not be an inevitable consequence of magma fragmentation. *Nature* **426**, 432–435 (2003).

30. Gardner, J. E., Llewellin, E. W., Watkins, J. M. & Befus, K. S. Formation of obsidian pyroclasts by sintering of ash particles in the volcanic conduit. *Earth Planet. Sci. Lett.* **459**, 252–263 (2017).

**Supplementary Information** is available in the online version of the paper.

**Acknowledgements** This study was supported by a European Research Council Advanced Grant to D.B.D. on 'Explosive volcanism in the Earth system: experimental insights' (EVOKES, grant number 247076). E.D. was supported by DFG grant ED 1757/1-1. We thank M. Kaliwoda and R. Hochleitner for Raman measurements at the Mineralogical State Collection Munich (SNSB). In addition, we thank H. W. Lohringer, D. Mueller, and A. Wimmer for assistance during the sample preparation, microprobe analyses and iron titration. Scientific discussions with M. Diez, D. Morgavi and C. M. P. De Campos were greatly appreciated. D.D.G. thanks C. Chelle-Michou and J. Lourenço for their assistance with R programming and H. Mader for comments and suggestions.

**Author Contributions** D.D.G. conceptualized the original idea, performed the low-temperature viscosity, calorimetric, Raman spectroscopic measurements, and Rhyolite-MELTS simulations. D.D.G., S.K. and S.W. synthesized the samples and performed the high-temperature viscosity measurements, wrote the original draft paper, finalized the figures, and compiled the database of chemical compositions. E.D. performed the magnetic measurements, wrote the associated text and finalized the figures. D.R.N, K.U.H. and D.B.D. advised on the optimal experimental methods and contributed to producing a final manuscript.

**Author Information** Reprints and permissions information is available at www.nature.com/reprints. The authors declare no competing financial interests. Readers are welcome to comment on the online version of the paper. Publisher's note: Springer Nature remains neutral with regard to jurisdictional claims in published maps and institutional affiliations. Correspondence and requests for materials should be addressed to D.D.G. (danilo.digenova@bristol.ac.uk).

**Reviewer Information** Nature thanks M. Manga and Y. Zhang for their contribution to the peer review of this work.
## METHODS

**Starting material preparation.** Experimental samples were synthesized to cover the range of K# and RAI of rhyolitic magmas from the Yellowstone volcanic system. Samples were mixed from powders of SiO₂, TiO₂, Al₂O₃, Fe₂O₃, MnO, MgO, CaCO₃, Na₂CO₃ and K₂CO₃. Loss of CO₂ during synthesis accounted for a final sample yield of 100 g. Powders were shaken before melting. Homogenization was performed twice. First, the sample was synthesized at 1,600 °C; the quenched glass was crushed and then milled in an agate mill. Final homogenization took place in a concentric cylinder viscometer by remelting and stirring each sample at the maximum possible rotation rate at 1,600 °C.

**Scanning electron microscope.** The samples were embedded in epoxy resin, polished and carbon-coated for electron microscopy. Backscattered electron image maps were collected at 5,000× and 10,000× magnification using a JEOL JSM IT300LV scanning electron microscope.

**Chemical analysis and iron oxidation state.** The chemical composition of the glasses was measured with a Cameca SX100 electron microprobe. Analyses were carried out at 15-kV acceleration voltage and, in order to minimise the alkali loss, a defocused beam (10 µm) and 5-nA beam current was used. Standards were wollastonite (Ca, Si), periclase (Mg), hematite (Fe), corundum (Al), natural orthoclase (K), and albite (Na). A matrix correction was performed according to ref. 31. The precision was better than 2.5% for all analysed elements. The chemical homogeneity of glasses was verified by performing around 25 chemical analyses for each sample.

About 100 mg of glass powder was used for the iron oxidation state determination. A redox titration using a standard solution containing potassium dichromate ($K_2Cr_2O_7$) was performed³² according to the following chemical reaction:

$$6\text{Fe}^{2+} + \text{Cr}_2\text{O}_7^{2-} + 14\text{H}^{+} \leftrightarrow 2\text{Cr}^{3+} + 6\text{Fe}^{3+} + 7\text{H}_2\text{O} \quad (2)$$

The titration allowed us to measure the amount of $Fe^{2+}$ in the glass and, in combination with electron microprobe analysis, the FeO/FeO₃ ratio was calculated for each sample. Titration was performed at 20 °C using a Metrohm Dosimat 665 automatic titrator with a 686 titroprocessor as the control unit. A basalt (BHVO-1, USGS standard, 8.58 wt% FeO) and a synthetic standard containing 18.8 wt% of FeO were used as standards. The standard deviation of the measurements was determined to be ±1.5%. The detailed procedure is reported in ref. 33.

**Concentric cylinder viscometry.** High-temperature viscometry was performed using a Brookfield DV-III+ viscometer head (full range of torque 0–0.7187 mN/m). The spindle used in these experiments was 42 mm long with 3.2 mm diameter and a 45° conical top and bottom to reduce edge effects³⁴⁻³⁶. The spindle and crucible were calibrated against the DGG1 standard glass, with an accurately known viscosity–temperature relationship. The precision of the viscosity determination³⁴ is ±3%. The temperature of the sample was calibrated to the furnace control thermocouple over the entire experimental temperature range using a separate platinum sheathed type-S thermocouple inserted directly into the sample.

The sample was held at the highest measurement temperature (1,590 °C) and stirred at the maximum achievable rotation rate for a minimum of 24 h before viscosity measurement in order to ensure chemical homogeneity, redox equilibration and the complete removal of potential bubbles from the melt. One study³⁷ reported complete redox equilibration in droplet-sized samples of iron-rich silicate melts to occur within less than 30 min, whereas others³⁸,³⁹ found that, for a sample of the same volume as used in this study but slightly higher iron contents than the ones used here, complete equilibration may take up to 24 h. We therefore consider that redox equilibrium was achieved during the super-liquidus melt viscosity determination, which lasted more than 24 h at atmospheric conditions.

The super-liquidus melt viscosity was measured at temperatures between 1,394 °C and 1,590 °C in steps of about 25°; see Extended Data Fig. 1. Each temperature step was maintained for at least 120 min and a stable torque reading was usually achieved after about 90 min, indicating thermal homogeneity. After the final measurement step, we re-measured the samples' viscosity at the highest imposed temperature to ensure measurement consistency and exclude drift in the measurement caused by mechanical effects in the measurement head or changes in sample composition (that is, loss of alkali elements at high temperature). In all cases, the torque values measured after re-occupying the highest experimental temperature stabilized with a reproducibility of about 1% of the measured value.

**Micropenetration.** Low-temperature viscosity measurements were performed using a Thermomechanical Analyzer (vertical dilatometer) TMA 402 F1/F3 Hyperion from Netzsch. The measuring system was thermally stabilized by water cooling. A turbo-molecular pump was used to reach the pressure of around $10^{-4}$ mbar necessary to create a vacuum in the chamber before measurement. Mass flow controllers for purge and protective gases were used to perform the viscosity measurements under a constant argon flow rate of 20 l min⁻¹. The temperature was calibrated with the known melting points of pure metals: indium (156.6 °C), zinc (419.5 °C), aluminium (660.3 °C) and silver (961.8 °C).

For each sample, we used two to three doubly polished disks (3 mm thick). The sample was heated up in two steps. Initially, a heating rate of 10 °C min⁻¹ was imposed up to a temperature 100 °C below the target temperature. Afterwards, for the last 100 °C, a rate of 3 °C min⁻¹ was employed in order to minimize the sample overshoot (±3 °C). At the target temperature, the sample was allowed to thermally equilibrate for 10 min. The viscosity measurement was performed at isothermal conditions applying a constant load, for approximately 10 min, to the sample surface via a push rod equipped with an iridium hemispherical indenter (radius $r = 1.3$ mm). The indentation was measured as a function of time via an inductive displacement transducer (resolution of 0.125 nm). The viscosity $\eta$ was calculated as follows⁴⁰,⁴¹:

$$\eta = \frac{0.1875Lt}{r^{0.5}\alpha^{1.5}} \quad (3)$$

where L is the load applied (in newtons), t the time (in seconds), r the indenter radius (in metres), and $\alpha$ the penetration (in metres). For each sample, we applied three different loads in order to reach approximately 100 µm of indentation during a single measurement. We applied 2.4 N for viscosity measurements at the lowest temperature, 0.6 N at the highest temperature, and 1.2 N for intermediate temperatures. The system was calibrated over a viscosity range between $10^8$ Pa s and $10^{12}$ Pa s with the standard glass DGG-1 of the Physikalisch-Technische Bundesanstalt. The error in the viscosity measurements was estimated to be ±0.05 logarithmic units.

**Glass transition measurements.** The glass transition temperature $T_g$ was measured using a Netzsch 404 Pegasus differential scanning calorimeter. The calorimeter was calibrated for temperature and sensitivity using the melting temperatures of pure metals. Approximately 50 mg of glass was loaded in a Pt crucible. Measurements were performed under a constant argon flow rate of 20 l min⁻¹. The sample was heated at 1 °C min⁻¹ from room temperature to 50 °C and kept at this temperature for approximately 30 min to achieve signal equilibrium in the differential scanning calorimeter. In order to structurally relax the glass, the sample was heated at 10 °C min⁻¹ up to approximately 50 °C above the peak of the glass transition. Afterwards, a cooling rate of 10 °C min⁻¹ and then a heating rate of 10 °C min⁻¹ were employed to measure the glass transition temperature at the peak ($T_{\text{peak},g}$, see Extended Data Fig. 5). Post-run optical analyses after each experiment confirmed the absence of crystallization of the samples (except for the L sample).

**Raman spectroscopy.** Raman spectra were acquired with an XploRa Raman System from Horiba. The instrument is equipped with an attenuated doubled Nd:YAG laser (532 nm). The laser power on the sample surface was measured to be about 3 mW through a 100× objective and a 1–2 µm² spot size. The instrument was calibrated using a silicon standard. Spectra were acquired using a grating of 1,800 grooves per millimetre, confocal hole diameter of 300 µm and slit width of 200 µm with an exposure time of 3 × 60 s. The backscattered Raman radiation was collected on a polished sample surface. The Raman signal was found to be optimal at a depth of 6 µm using a motor on the Z axis and, consequently, the spectra were collected at the same depth. Prior the Raman spectra acquisition, the samples were stored at 100 °C in an oven to avoid water absorption on the surface. The samples were found to be nominally water-free.

**Magnetic analyses.** Hysteresis loops were generated using a Lake Shore 7400 vibrating sample magnetometer. The magnetization of the specimens was measured by applying a field ranging between –1 T and +1 T to six specimens obtained from samples F and J and a field varying between –1.5 T and +1.5 T to ten specimens obtained from sample L.

**Rheological agpaitic index.** Rhyolites are commonly classified according to the agpaitic index (AI), which is defined as the ratio of the molar sum of alkali oxides over aluminium oxide—($Na_2O + K_2O$)/Al₂O₃—where AI > 1, AI < 1 and AI = 1 define peralkaline, metaluminous and peraluminous melts, respectively. Peralkaline melts are known to be much less viscous than metaluminous and peraluminous melts⁴²⁻⁴⁴. These studies show that such rare compositions (for example, comendite and pantellerite) have a host of other physicochemical properties that deviate substantially from those of calcalkaline rhyolites (lower viscosities, higher volatile solubilities, higher diffusivities, higher surface tensions, different redox state, and so on)⁴⁴. Such compositions are, therefore, not included in our analysis and we focus on the transport properties and eruptive styles of calcalkaline rhyolites, which volumetrically dominate rhyolitic volcanism on Earth.

We introduce a rheological agpaitic index (termed RAI; see equation (1)), which considers the molar sum of alkali, alkaline earth oxides, and FeO over the sum of Fe₂O₃ and Al₂O₃ to classify calcalkaline rhyolitic magmas from a melt structural and rheological point of view. Using this new classification, we can distinguish samples characterized by: (1) RAI ≤ 1, where not tetrahedrally coordinated cations (alkali, alkaline earths, and Fe²⁺) behave as charge balancers of the tetrahedrally coordinated cations (Al³⁺ and Fe³⁺) and (2) RAI > 1, where the excess
not tetrahedrally coordinated cations start to behave as network modifiers. In the latter case, the excess of not tetrahedrally coordinated cations over tetrahedrally coordinated ones depolymerizes the melt structure and decreases the melt viscosity. **Source of error in RAI determination.** Two sources of error may arise in the calculation of RAI:

(1) Analytical uncertainty: common uncertainties of electron microprobe analyses range around 2%–3% for most major elements and up to 10% for sodium depending on the measurement conditions. When considering, for example, the eruption of Chaitén⁴⁵, for which the calculated RAI lies close to the rheological discontinuity (RAI = 1), the errors arising from analytical uncertainty may produce variations in RAI of up to 0.04, resulting in a RAI value just over or just under 1.

(2) Assumptions about the iron oxidation state: as iron changes its melt structural role depending on its oxidation state, assumptions on the $Fe^{2+}/Fe^{3+}$ ratio in a melt may change the calculated RAI. The magnitude of this change is dependent on the absolute Fe content and may reach up to 0.04 RAI units for a 40% change in $Fe^{2+}/Fe^{3+}$ ratio in melts with relatively high iron content (independent of the analytical uncertainty). For the case of Chaitén⁴⁵, for example, the calculated RAI shifts from 0.93 to 1.1 by assuming a 30/70 and 70/30 percentage ratio of $Fe^{2+}/Fe^{3+}$, respectively.

The correct assessment of such borderline cases thus depends on a detailed knowledge of the analytical uncertainty and $f_{O_2}$. Although the theoretical threshold for the melt structural point of change is at RAI=1, the above-mentioned analytical contributions may shift certain literature data into the field of opposite eruptive response.

Further, it is important to note that the chemical analyses used to apply RAI must be representative of the melt compositions (that is, the matrix glass) and not whole-rock analyses of partially crystalline volcanic rocks. This is because the viscous response of the magma on volcanic timescales occurs only in the melt. Since the changes in composition that influence the melts' RAI parameter are small, using whole-rock analyses that include varying percentages of crystal cargo or X-ray fluorescence data from crystal-poor samples may push the composition to a RAI value that is not representative of the melts' eruptive style.

**Comparison of the viscosity measurements with existing models.** Our data show that current viscosity models (HZ¹¹ and GRD¹²) for multicomponent systems do not capture the changes in viscosity caused by K# and RAI (Fig. 1). In Extended Data Fig. 2 we report a comparison between our measurements and predictions using the most widely used models¹¹,¹². Both models are quasi-insensitive to changes in K# (see samples E and A, respectively K₂O-rich and K₂O-poor), showing a predicted viscosity of about 10¹⁰ Pa s, while the measured viscosities span two orders of magnitude (Fig. 1). Similarly, the measured abrupt drop in viscosity due to the change in RAI from <1 (sample D) to >1 (sample I) is not predicted.

**Iron oxide nanolite identification and their formation.** Scanning electron microscope analysis was performed to detect the possible presence of crystals. Images collected at 5,000× and 10,000× magnifications of the sample before and after the viscosity measurement did not reveal the presence of any crystalline phases. Contrary to the scanning electron microscope analysis, Raman spectra (Extended Data Fig. 3) showed the presence of a crystalline iron oxide peak. Although the Raman spectrum of sample L preserved the typical Raman features of rhyolitic glasses⁴⁶⁻⁴⁸, the overall Raman signature is considerably affected by a well defined peak located at about 670 cm⁻¹ attributed to iron oxide (RUFF Raman spectra database⁵¹; http://rruff.info/). This peak was not detected in the Raman spectra of samples F and J (Extended Data Fig. 3).

Magnetic hysteresis analyses of the samples are in agreement with the results from Raman spectroscopy. The loops obtained from samples F and J are reversible and linear with a positive (paramagnetic) susceptibility, which varies in the range 2.29–5.35 × 10⁻² A m² kg⁻¹ T⁻¹ (Extended Data Fig. 4a). Paramagnetic behaviour indicates the absence of electronic exchange interaction typical of ferromagnetic minerals such as common iron oxides⁴⁹. The ten specimens obtained from sample L are characterized by nonlinear hysteresis behaviour (Extended Data Fig. 4b). All loops are reversible and closed, indicating the absence of a remanent magnetization ($M_r$ that is, the remanence left after reducing the applied field to zero). This is typical of superparamagnetic particles⁵⁰,⁵²⁻⁵⁴. Superparamagnetism is a state that characterizes ferromagnetic particles with a relaxation time short enough (tens of seconds) to acquire a zero net magnetization within laboratory analysis time. For equidimensional magnetite, the size threshold from the superparamagnetic to the stable single-domain state (that is, capable of acquiring a stable $M_r$) has been estimated previously⁵⁰,⁵³,⁵⁵ to be about 20–30 nm. The superparamagnetic hysteresis can be described by the Langevin function:

$$ \frac{M}{M_s} = N \left( \coth \gamma - \frac{1}{\gamma} \right) \qquad (4) $$

where M is the magnetization, $M_s$ is the saturation magnetization⁵⁶, and N is the number of moments. The term $\gamma$ is defined as:

$$ \gamma = \frac{M_s Bv}{kT} \qquad (5) $$

where B is the applied field, $v$ is the particle volume, $k$ is Boltzmann's constant, and T is the absolute temperature. For superparamagnetic particles for which the Langevin function is applicable⁵³, B90 (that is, the value at which $0.9M_s$ is reached) occurs at $\gamma \approx 10$. In sample L, after correction for a paramagnetic slope of the loops probably deriving from iron ions dispersed within the matrix, B90 has an average value of $0.63 \pm 0.02 T$ (Extended Data Fig. 4c). Assuming cubic particles of $v = d^3$, the grain size of the particles can be estimated from equation (5) as follows:

$$ d = 3 \sqrt{\frac{10kT}{B90 \times M_s}} \qquad (6) $$

This relationship is represented in Extended Data Fig. 4d. For sample L, it results in an average $d = 5.2 \pm 0.05$ nm. This result allows us to constrain the size of the magnetite grains in the specimen between about 5 nm and about 30 nm, the maximum estimated threshold size for the superparamagnetic–single domain state change.

Nanolite particles formed during cooling after high-temperature viscosity measurements (Extended Data Fig. 1). In the high-temperature regime, the viscosity decreases with decreasing the $SiO_2$ content and, therefore, the L sample ($SiO_2$ content 71.22 wt%) is the least viscous. In the low-temperature regime, however, the L sample displays unexpectedly high viscosity values. We can, therefore, constrain the nanolite formation regime to lie between 1,420 °C (lowest temperature investigated during high-temperature viscosity measurements) and the glass transition interval during melt cooling. Particularly, the decrease of temperature oxidized the melt by increasing the $Fe^{3+}/Fe^{2+}$ ratio owing to the temperature dependence of $f_{O_2}$. This process decreases the total amount of network modifier cations and transforms them into charge-balancer cations necessary to stabilize the $Fe^{3+}$ in four-fold coordination (network former). The resulting double-feedback effect lies at the heart of this rapid and drastic change in melt structure and, therewith, viscosity. Once the charge-balancer cations are consumed, iron-oxide nanoparticles precipitate from the melt. As a result, the degree of polymerization of the melt structure and viscosity increases. According to the L-sample chemistry (Extended Data Table 1) and assuming the entire extraction of magnetite⁵⁷, the residual melt would be depleted in $FeO_{tot}$ and the residual iron (0.66 mol%) would be present as $Fe_2O_3$.

Calorimetric analysis of the sample revealed two endothermic peaks in the glass transition interval (Extended Data Fig. 5) locked in during quenching of the L sample from high temperature. This is an unusual phenomenon as only one endothermic peak would be expected in the glass transition interval⁵⁸,⁵⁹. Two studies reported the same observation for a basaltic–andesitic scoria produced by a violent Strombolian eruption at Mt Yasur Volcano in Vanuatu²⁷ and a series of basaltic melts characterized by different oxygen fugacities⁶⁰. The authors attributed these peaks to the coexistence of (1) two chemically invariant glasses having two redox states and/or (2) two glasses having different iron content and chemical composition (that is, Fe-rich and Fe-poor). We propose, on the basis of the experimental data presented here, that sample L contained two distinct liquid domains with different iron contents and probably also different redox states. The first domain, nanolite-free and less polymerized, preserving the total amount of iron in the amorphous structure, showed the lowest $T_g^{peak}$ (at 686.6 °C). The second, more polymerized domain, in which iron was removed from the melt structure to form nanolites, showed a higher $T_g^{peak}$ (732.1 °C). With the second heat treatment, the nanolite-free domain vanished, as confirmed by the disappearance of its $T_g^{peak}$. When performing micro-penetration viscosity measurements on the L sample, well above the glass transition interval, the relatively slow heating rate (10 °C min⁻¹) allows the melt to oxidize further, and hence the nanolite content increased further, leading to an increase in the melt polymerization and viscosity (Fig. 2).

We compared the $T_g^{peak}$ of sample L with the $T_g^{peak}$ of the nanolite-free samples F ($FeO_{tot}$ 1.55 wt%) and J ($FeO_{tot}$ 3.31 wt%); see Extended Data Fig. 5. In the glass transition interval only one well defined peak is observed confirming the homogeneity of these samples. As expected, $T_g^{peak}$ decreases with decreasing $SiO_2$ content: 879.8 °C and 757.6 °C for the F and J samples, respectively.

To estimate the viscosity of the nanolite-free domain of sample L and therefore evaluate the effect on nanolite formation on the melt viscosity, we calculated the viscosity at $T_g^{peak}$ (Extended Data Table 1) for all the samples investigated in this study (Extended Data Fig. 7). To do so, in the investigated interval of temperature, we fitted Arrhenian models to the viscosity data of each sample. Afterwards, following the approach of ref. 61, we used the relationship reported in equation (7) between the viscosity and the glass transition temperature at a given heating or cooling rate, which is considered to be independent of the melt composition:

$$ \log(\eta_{T_{glass}}) = K - \log(q) \qquad (7) $$
where $\eta_{Tglass}$ and q represent the viscosity at the glass transition and the cooling or heating rate (in °C s⁻¹), respectively. K is the shift factor, which we calculated to be 10.61 ± 0.29 (excluding the nanolite-bearing sample L). This is in agreement with the K = 10.65 ± 0.28 value provided for calcalkaline rhyolites⁶¹. We can now calculate the expected viscosity (10¹¹.³⁹ Pa s) of sample L at $T_{\text{peak}}$ (L-DTA sample in Fig. 2). Considering the temperature dependence of the viscosity of sample J, we estimated an increase of viscosity of about 10².⁵ Pa s, attributable to the presence of nanolite particles.

**Rhyolite-MELTS modelling.** A rhyolitic liquid with equal molar amounts of K₂O and Na₂O (K# = 0.5) of 1 was chosen as the starting composition. We modelled the paths of equilibrium and fractional crystallization, at two pressures P (0.5 and 1.5kbar) for two different $f_{O_2}$ values (Fig. 3a), from 1,050 °C to 750 °C.

The model results show that equilibrium melt crystallization strongly decreases RAI (Fig. 3a) while increasing K# (Fig. 3b). The temperature dependence of $f_{O_2}$ during the simulations varied according to the QFM and QFM + 2 buffer reactions in order to mimic the fugacities commonly assumed for the crust⁶². This allows to constrain the evolution of the solid and liquid composition and the phase assemblage evolution with $T-P-f_{O_2}$ and calculate the evolution of RAI and K# with temperature (Fig. 3a and b).

At a fixed temperature, with decreasing $f_{O_2}$ of the system (from QFM + 2 to QFM buffer) RAI increases, surpassing the critical rheological threshold of RAI = 1 (Fig. 3a). $f_{O_2}$ does not much affect K#. However, pressure and water degassing play a major part in increasing K#, particularly at low temperatures (between 750 °C and 800 °C) (Fig. 3b). When modelling fractional crystallization K# increases further (Supplementary Table 1).

In summary, a low degree of crystallization and, especially, reduced $f_{O_2}$ conditions near the QFM buffer favour the genesis of low-viscosity rhyolites (RAI > 1 and low K#). By contrast, RAI decreases and K# increases as crystallization proceeds, promoting the formation of a highly polymerized melt. These structural changes can vary the melt viscosity by up to two orders of magnitude at constant environmental parameters (T, P, xH₂O).

This is due to the complex and simultaneous effects of $T-P-f_{O_2}$ on the residual liquid- and solid-phase assemblage composition (Extended Data Table 1). The crystalline phases that influence the residual melt composition and viscosity the most strongly are magnetite and orthopyroxene, followed by albite and sanidine. With decreasing temperature, the magnetite and orthopyroxene precipitation, together with the oxidation of the iron in the melt (at fixed oxygen buffer), strongly increase the $Fe^{3+}/Fe^{2+}$ ratio of the residual melt and consequently decrease RAI (Fig. 3a). The crystallization of feldspars (mainly albite) not only decreases RAI (residual melt is enriched in Al₂O₃) but also greatly increases K# via the K₂O content enrichment of the melt (Fig. 3b). This results in a double effect of raising the melt viscosity and hence also the explosive potential of the magma. We note that the above-discussed changes apply to the residual melt viscosity only, and any effect of increased solid fraction on the effective viscosity of the magmatic mixture is not discussed here. Crystallization (that is, increasing solid fraction) will act to increase the effective viscosity of the magmatic suspension further, in addition to the increase in residual melt viscosity⁶³,⁶⁴.

Melts characterized by high K# may also be produced by open-system processes owing to an additional input of K₂O. In fact, host rock assimilation is frequently suggested to be of the order of tens of per cent¹⁸, which greatly enhances the potential for changes in RAI and K#. Irrespective of the role of open-system processes, however, our simulations demonstrate that crystallization, changes in $f_{O_2}$ or variations in pressure may cause critical variations in melt chemistry, leading the melt to cross the experimentally observed rheological transition, affecting its fragmentation and welding behaviour²⁹,³⁰,⁶⁵.

**Global database of composition of rhyolitic systems.** The chemical analyses used to calculate K# and RAI in Fig. 4 were selected from 64 publications¹⁰,¹⁶,¹⁷,²⁰,⁶⁶⁻¹²⁴ reporting chemical compositions of glass matrix analyses of rhyolitic products erupted effusively or explosively across the planet. For visual simplicity, the RAI and K# values plotted in Fig. 4b are averaged values of the database contents for each melt.

**Data availability.** The data that support the findings of this study are available from the corresponding author upon reasonable request.

31. Pouchou, J.-L. & Pichoir, F. in *Electron Probe Quantitation* SE-4 (eds Heinrich, K. F. J. & Newbury, D.) 31–75 (Springer, 1991).

32. Shapiro, L. M. & Brannock, W. W. Rapid analysis of silicate rocks. *Geol. Surv. Bull.* C 1036, 295–318 (1956).

33. Giuli, G. et al. XAS determination of the Fe local environment and oxidation state in phonolite glasses. *Am. Mineral.* **96**, 631–636 (2011).

34. Dingwell, D. B. Viscosity–temperature relationships in the system Na₂Si₂O₅–Na₄Al₂O₅. *Geochim. Cosmochim. Acta* **50**, 1261–1265 (1986).

35. Dingwell, D. B. Shear viscosities of ferrosilicate liquids. *Am. Mineral.* **74**, 1038–1044 (1989).

36. Dingwell, D. B. & Virgo, D. Viscosities of melts in the Na₂O–FeO–Fe₂O₃–SiO₂ system and factors controlling relative viscosities of fully polymerized silicate melts. *Geochim. Cosmochim. Acta* **52**, 395–403 (1988).

37. Mysen, B. O. & Virgo, D. Influence of pressure, temperature, and bulk composition on melt structures in the system NaAlSi₂O₆–NaFe₃ + Si₂O₆. *Am. J. Sci.* **278**, 1307–1322 (1978).

38. Dingwell, D. B. & Virgo, D. The effect of oxidation state on the viscosity of melts in the system Na₂O–FeO–Fe₂O₃–SiO₂. *Geochim. Cosmochim. Acta* **51**, 195–205 (1987).

39. Dingwell, D. B. Redox viscometry of some Fe-bearing silicate melts. *Am. Mineral.* **76**, 1560–1562 (1991).

40. Pocklington, H. C. Rough measurement of high viscosities. *Math. Proc. Camb. Phil. Soc.* **36**, 507–508 (1940).

41. Tobolsky, A. V. & Taylor, R. B. Viscoelastic properties of a simple organic glass. *J. Phys. Chem.* **67**, 2439–2442 (1963).

42. Scarfe, C. M. Melt at one atmosphere viscosity of a pantellerite. *Can. Mineral.* **15**, 185–189 (1977).

43. Stevenson, R. J., Bagdassarov, N. S., Dingwell, D. B. & Romano, C. The influence of trace amounts of water on the viscosity of rhyolites. *Bull. Volcanol.* **60**, 89–97 (1998).

44. Di Genova, D. et al. The rheology of peralkaline rhyolites from Pantelleria Island. *J. Volcanol. Geotherm. Res.* **249**, 201–216 (2013).

45. Watt, S. F. L., Pyle, D. M. & Mather, T. A. Evidence of mid- to late-Holocene explosive rhyolitic eruptions from Chaitén volcano, Chile. *Andean Geol.* **40**, 216–226 (2013).

46. Di Genova, D. et al. Approximate chemical analysis of volcanic glasses using Raman spectroscopy. *J. Raman Spectrosc.* **46**, 1235–1244 (2015).

47. Di Genova, D. et al. Raman spectra of Martian glass analogues: a tool to approximate their chemical composition. *J. Geophys. Res. Planets* **121**, 740–752 (2016).

48. Di Genova, D., Hess, K.-U., Chevrel, M. O. & Dingwell, D. B. Models for the estimation of $Fe^{3+}/Fe_{tot}$ ratio in terrestrial and extra-terrestrial alkali- and iron-rich silicate glasses using Raman spectroscopy. *Am. Mineral.* **101**, 943–952 (2016).

49. Tauxe, L. *Essentials of Paleomagnetism* Ch. 3, 33–46 (Univ. California Press, 2010).

50. Dunlop, D. J. Superparamagnetic and single-domain threshold sizes in magnetite. *J. Geophys. Res.* **78**, 1780–1793 (1973).

51. Lafuente, B., Downs, R. T., Yang, H. & Stone, N. The power of databases: the RRUFF project. In *Highlights in Mineralogical Crystallography* (eds Armbruster, T. & Danisi, R. M.) 1–30 (De Gruyter, 2015).

52. Dunlop, D. J. Theory and application of the Day plot (Mrs/Ms versus Hcr/Hc). *1. Theoretical curves and tests using titanomagnetite data*. *J. Geophys. Res. Solid Earth* **107**, 1–22 (2002).

53. Tauxe, L., Mullender, T. A. T. & Pick, T. Potbellies, wasp waists, and superparamagnetism in magnetic hysteresis. *J. Geophys. Res.* **101**, 571–583 (1996).

54. Tauxe, L., Bertram, H. N. & Seberino, C. Physical interpretation of hysteresis loops: micromagnetic modelling of fine particle magnetite. *Geochem. Geophys. Geosyst.* **3**, 1–22 (2002).

55. Butler, R. F. & Banerjee, S. K. Theoretical single-domain grain size range in magnetite and titanomagnetite. *J. Geophys. Res.* **80**, 4049 (1975).

56. Pullaiah, G., Irving, E., Buchanan, K. L. & Dunlop, D. J. Magnetization changes caused by burial and uplift. *Earth Planet. Sci. Lett.* **28**, 133–143 (1975).

57. Liebske, C., Behrens, H., Holtz, F. & Lange, R. A. The influence of pressure and composition on the viscosity of andesitic melts. *Geochim. Cosmochim. Acta* **67**, 473–485 (2003).

58. Di Genova, D., Romano, C., Giordano, D. & Alletti, M. Heat capacity, configurational heat capacity and fragility of hydrous magmas. *Geochim. Cosmochim. Acta* **142**, 314–333 (2014).

59. Bouhifid, M. A., Whittington, A., Withers, C. & Richet, P. Heat capacities of hydrous silicate glasses and liquids. *Chem. Geol.* **346**, 125–134 (2013).

60. Chevrel, M. O., Baratoux, D., Hess, K.-U. & Dingwell, D. B. Viscous flow behavior of tholeiitic and alkaline Fe-rich martian basalts. *Geochim. Cosmochim. Acta* **124**, 348–365 (2014).

61. Stevenson, R. J., Dingwell, D. B., Webb, S. L. & Bagdassarov, N. S. The equivalence of enthalpy and shear stress relaxation in rhyolitic obsidians and quantification of the liquid-glass transition in volcanic processes. *J. Volcanol. Geotherm. Res.* **68**, 297–306 (1995).

62. Philpotts, A. R & Ague, J. J. *Principles of Igneous and Metamorphic Petrology* Ch. 11 (Cambridge Univ. Press, 2009).

63. Kolzenburg, S., Giordano, D., Cimarelli, C. & Dingwell, D. B. *In situ* thermal characterization of cooling/crystallizing lavas during rheology measurements and implications for lava flow emplacement. *Geochim. Cosmochim. Acta* **195**, 244–258 (2016).

64. Kolzenburg, S., Giordano, D., Thordarson, T., Höskuldsson, A. & Dingwell, D. B. The rheological evolution of the 2014/2015 eruption at Holuhraun, central Iceland. *Bull. Volcanol.* **79**, 45 (2017).

65. Kolzenburg, S. & Russell, J. K. Welding of pyroclastic conduit infill: a mechanism for cyclical explosive eruptions. *J. Geophys. Res. Solid Earth* **119**, 5305–5323 (2014).

66. Perkins, M. E., Nash, W. P., Brown, F. H. & Fleck, R. J. Fallout tuffs of Trapper Creek, Idaho—a record of Miocene explosive volcanism in the Snake River Plain volcanic province. *Geol. Soc. Am. Bull.* **107**, 1484–1506 (1995).

67. Perkins, M. E. & Nash, B. P. Explosive silicic volcanism of the Yellowstone hotspot: the ash fall tuff record. *Geol. Soc. Am. Bull.* **114**, 367–381 (2002).
68. Nash, B. P. & Perkins, M. E. Neogene fallout tuffs from the Yellowstone hotspot in the Columbia Plateau region, Oregon, Washington and Idaho, USA. *PLoS One* **7**, 1–13 (2012).

69. Girard, G. & Stix, J. Magma recharge and crystal mush rejuvenation associated with early post-collapse Upper Basin Member rhyolites, Yellowstone caldera, Wyoming. *J. Petrol.* **50**, 2095–2125 (2009).

70. Ellis, B. S. et al. Groundmass crystallisation and cooling rates of lava-like ignimbrites: the Grey's Landing ignimbrite, southern Idaho, USA. *Bull. Volcanol.* **77**, 87 (2015).

71. Ellis, B. S. et al. Petrologic constraints on the development of a large-volume, high temperature, silicic magma system: The Twin Falls eruptive centre, central Snake River Plain. *Lithos* **120**, 475–489 (2010).

72. Rowe, M. C., Ellis, B. S. & Lindeberg, A. Quantifying crystallization and devitrification of rhyolites by means of X-ray diffraction and electron microprobe analysis. *Am. Mineral.* **97**, 1685–1699 (2012).

73. Cathey, H. E. & Nash, B. P. The Cougar Point tuff: implications for thermochemical zonation and longevity of high-temperature, large-volume silicic magmas of the Miocene Yellowstone hotspot. *J. Petrol.* **45**, 27–58 (2004).

74. Bolte, T., Holtz, F., Almeev, R. & Nash, B. P. The Blacktail Creek tuff: an analytical and experimental study of rhyolites from the Heise volcanic field, Yellowstone hotspot system. *Contrib. Mineral. Petrol.* **169**, 15 (2015).

75. Clay, P. L. et al. Textural characterization, major and volatile element quantification and Ar-Ar systematics of spherulites in the Rocche Rosse obsidian flow, Lipari, Aeolian Islands: a temperature continuum growth model. *Contrib. Mineral. Petrol.* **165**, 373–395 (2013).

76. de Rosa, R. & Sheridan, M. F. Evidence for magma mixing in the surge deposits of the Monte Guardia sequence, Lipari. *J. Volcanol. Geotherm. Res.* **17**, 313–328 (1983).

77. de Rosa, R., Donato, P., Gioncada, A., Masetti, M. & Santacroze, R. The Monte Guardia eruption (Lipari, Aeolian Islands): an example of a reversely zoned magma mixing sequence. *Bull. Volcanol.* **65**, 530–543 (2003).

78. Pallister, J. S. et al. The Chaitén rhyolite lava dome: eruption sequence, lava dome volumes, rapid effusion rates and source of the rhyolite magma. *Andean Geol.* **40**, 277–294 (2013).

79. Castro, J. M. & Dingwell, D. B. Rapid ascent of rhyolitic magma at Chaitén volcano, Chile. *Nature* **461**, 780–783 (2009).

80. Folkes, C. B., de Silva, S. L., Wright, H. M. & Cas, R. Geochemical homogeneity of a long-lived, large silicic system; evidence from the Cerro Galán caldera, NW Argentina. *Bull. Volcanol.* **73**, 1455–1486 (2011).

81. Piochi, M. et al. Constraining the recent plumbing system of Volcano (Aeolian Arc, Italy) by textural, petrological, and fractal analysis: the 1739 AD Pietro Cotte lava flow. *Geochem. Geophys. Geosyst.* **10**, Q01009 (2009).

82. Befus, K. S., Zinke, R. W., Jordan, J. S., Manga, M. & Gardner, J. E. Pre-eruptive storage conditions and eruption dynamics of a small rhyolite dome: Douglas Knob, Yellowstone volcanic field, USA. *Bull. Volcanol.* **76**, 808 (2014).

83. Castro, J. M. et al. Storage and eruption of near-liquidus rhyolite magma at Cordon Caulle, Chile. *Bull. Volcanol.* **75**, 702 (2013).

84. Alloway, B. V., Pearce, N. J. G., Villarosa, G., Outes, V. & Moreno, P. I. Multiple melt bodies fed the AD 2011 eruption of Puyehue-Cordón Caulle, Chile. *Sci. Rep.* **5**, 17589 (2015).

85. Fytikas, M. et al. Volcanology and petrology of volcanic products from the island of Milos and neighbouring islets. *J. Volcanol. Geotherm. Res.* **28**, 297–317 (1986).

86. Gertisser, R., Preece, K. & Keller, J. The Plinian Lower Pumice 2 eruption, Santorini, Greece: magma evolution and volatile behaviour. *J. Volcanol. Geotherm. Res.* **186**, 387–406 (2009).

87. Michaud, V., Clocchiatti, R. & Sbrana, S. The Minoan and post-Minoan eruptions, Santorini (Greece), in the light of melt inclusions: chlorine and sulphur behaviour. *J. Volcanol. Geotherm. Res.* **99**, 195–214 (2000).

88. Heiken, G. Plinian-type eruptions in the Medicine Lake Highland, California, and the nature of the underlying magma. *J. Volcanol. Geotherm. Res.* **4**, 375–402 (1978).

89. Eichelberger, J. C. Origin of andesite and dacite: evidence of mixing at Glass Mountain in California and at other circum-Pacific volcanoes. *Bull. Geol. Soc. Am.* **86**, 1381–1391 (1975).

90. Foresta Martin, F. et al. New insights into the provenance of the obsidian fragments of the island of Ustica (Palermo, Sicily). *Archaeometry* **59**, 435–454 (2016).

91. Rosen, S. A., Tykot, R. H. & Gottesman, M. Long distance trinket trade: Early Bronze Age obsidian from the Negev. *J. Archaeol. Sci.* **32**, 775–784 (2005).

92. Shackley, M. S. Geologic origin of the source of Bearhead Rhyolite (Paliza Canyon) obsidian, Jemez Mountains, Northern New Mexico. *New Mex. Geol.* **38**, 52–65 (2016).

93. Ulrich, T. & Kamber, B. S. Natural obsidian glass as an external accuracy reference material in laser ablation-inductively coupled plasma-mass spectrometry. *Geostand. Geoanal. Res.* **37**, 169–188 (2013).

94. Pe-Piper, G. & Moulton, B. Magma evolution in the Pliocene-Pleistocene succession of Kos, South Aegean arc (Greece). *Lithos* **106**, 110–124 (2008).

95. Dunbar, N. W., Hervig, R. L. & Kyle, P. R. Determination of pre-eruptive H2O, F and Cl contents of silicic magmas using melt inclusions: examples from Taupo volcanic center, New Zealand. *Bull. Volcanol.* **51**, 177–184 (1989).

96. Stokes, S. & Lowe, D. J. Discriminant function analysis of late Quaternary tephras from five volcanoes in New Zealand using glass late major element chemistry. *Quat. Res.* **30**, 270–283 (1992).

97. Bindeman, I. N. & Valley, J. W. Low-$\delta^{18}$O rhyolites from Yellowstone: magmatic evolution based on analyses of zircons and individual phenocrysts. *J. Petrol.* **42**, 1491–1517 (2001).

98. Chamberlain, K. J., Morgan, D. J. & Wilson, C. J. N. Timescales of mixing and mobilisation in the Bishop Tuff magma body: perspectives from diffusion chronometry. *Contrib. Mineral. Petrol.* **168**, 1034 (2014).

99. Hildreth, W. & Wilson, C. J. N. Compositional zoning of the bishop tuff. *J. Petrol.* **48**, 951–999 (2007).

100. Watt, S. F. L., Pyle, D. M., Mather, T. A., Martin, R. S. & Matthews, N. E. Fallout and distribution of volcanic ash over Argentina following the May 2008 explosive eruption of Chaitén, Chile. *J. Geophys. Res.* **114**, B04207 (2009).

101. Hammer, J. E., Cashman, K. V., Hoblitt, R. P. & Newman, S. Degassing and microlite crystallization during pre-climatic events of the 1991 eruption of Mt Pinatubo, Philippines. *Bull. Volcanol.* **60**, 355–380 (1999).

102. Leonard, G. S., Cole, J. W., Nairn, I. A. & Self, S. Basalt triggering of the c. AD 1305 Kaharoa rhyolite eruption, Tarawera volcanic complex, New Zealand. *J. Volcanol. Geotherm. Res.* **115**, 461–486 (2002).

103. Arce, J. L., Cervantes, K. E., Macías, J. L. & Mora, J. C. The 12.1 ka Middle Toluca Pumice: a dacitic Plinian-subplinian eruption of Nevado de Toluca in Central Mexico. *J. Volcanol. Geotherm. Res.* **147**, 125–143 (2005).

104. Scaillet, B. & Evans, B. W. The 15 June 1991 eruption of Mount Pinatubo. I. Phase equilibria and pre-eruption P-T-fO$_2$-fH$_2$O conditions of the dacite magma. *J. Petrol.* **40**, 381–411 (1999).

105. Walter, R. C., Hart, W. K. & Westgate, J. A. Petrogenesis of a basalt-rhyolite tephra from the west-central Afar, Ethiopia. *Contrib. Mineral. Petrol.* **95**, 462–480 (1987).

106. Selbekk, R. S. & Trønnes, R. G. The 1362 AD Öræfajökull eruption, Iceland: petrology and geochemistry of large-volume homogeneous rhyolite. *J. Volcanol. Geotherm. Res.* **160**, 42–58 (2007).

107. Palais, J. M. & Sigurdsson, H. Petrologic evidence of volatile emissions from major historic and pre-historic volcanic eruptions. *Understanding Clim. Change* **52**, 31–53 (1989).

108. Sharma, K., Self, S., Blake, S., Thordarson, T. & Larsen, G. The AD 1362 Öræfajökull eruption, S.E. Iceland: physical geochemistry and volatile release. *J. Volcanol. Geotherm. Res.* **178**, 719–739 (2008).

109. Macdonald, R. et al. The 1875 eruption of Askja volcano, Iceland: combined fractional crystallization and selective contamination in the generation of rhyolitic magma. *Mineral. Mag.* **51**, 183–202 (1987).

110. Gunnarsson, B., Marsh, B. D. & Taylor, H. P. Generation of Icelandic rhyolites: silicic lavas from the Torfajökull central volcano. *J. Volcanol. Geotherm. Res.* **83**, 1–45 (1998).

111. Tuffen, H. & Castro, J. M. The emplacement of an obsidian dyke through thin ice: Hrafninnuhryggur, Krafla Iceland. *J. Volcanol. Geotherm. Res.* **185**, 352–366 (2009).

112. Macías, J. L. et al. A 550-year-old Plinian eruption at El Chichón Volcano, Chiapas, Mexico: explosive volcanism linked to reheating of the magma reservoir. *J. Geophys. Res.* **B 108**, ECV3-1-18 (2003).

113. Sparks, R. S. J., Wilson, L. & Sigurdsson, H. The pyroclastic deposits of the 1875 eruption of Askja, Iceland. *Phil. Trans. R. Soc. Lond. A* **299**, 242–272 (1981).

114. Baker, P. E., Buckley, F. & Holland, J. G. Petrology and chemistry of Easter Island. *Contrib. Mineral. Petrol.* **44**, 85–100 (1974).

115. Edmonds, M., Pyle, D. & Oppenheimer, C. A model for degassing at the Soufriere Hill Volcano, Montserrat, West Indies, based on geochemical data. *Earth Planet. Sci. Lett.* **186**, 159–173 (2001).

116. Rango, T., Colombani, N., Mastrocicco, M., Bianchini, G. & Beccaluva, L. Column elution experiments on volcanic ash: geochemical implications for the main Ethiopian rift waters. *Wat. Air Soil Pollut.* **208**, 221–233 (2010).

117. Manley, C. R. Morphology and maturation of melt inclusions in quartz phenocrysts from the Badlands rhyolite lava flow, southwestern Idaho. *Am. Mineral.* **81**, 158–168 (1996).

118. Higgins, M. W. Petrology of Newberry volcano, Central Oregon. *Bull. Geol. Soc. Am.* **84**, 455–488 (1973).

119. Stevenson, R. J., Hodder, A. P. W. & Briggs, R. M. Rheological estimates of rhyolite lava flows from the Okataina volcanic centre, New Zealand. *N. Z. J. Geol. Geophys.* **37**, 211 (1994).

120. Kolzenburg, S., Russell, J. K. & Kennedy, L. A. Energetics of glass fragmentation: experiments on synthetic and natural glasses. *Geochem. Geophys. Geosyst.* **14**, 4936–4951 (2013).

121. Albert, P. G. et al. Marine-continental tephra correlations: Volcanic glass geochemistry from the Marsili Basin and the Aeolian Islands, Southern Tyrrhenian Sea, Italy. *J. Volcanol. Geotherm. Res.* **229**, 74–94 (2012).

122. Singer, B. S. et al. Eruptive history, geochronology, and magmatic evolution of the Puyehue-Cordón Caulle volcanic complex, Chile. *Bull. Geol. Soc. Am.* **120**, 599–618 (2008).

123. Kuehn, S. C. & Foit, F. F. Jr. Silicic tephras of Newberry Volcano. In *What's New at Newberry Volcano, Oregon: Guidebook for the Friends of the Pleistocene Eighth Annual Pacific Northwest Cell Field Trip* (eds Jensen, R. A. & Chitwood, L. A.) 135–163 (2000).

124. Schipper, C. I. et al. Cristobalite in the 2011–2012 Cordón Caulle eruption (Chile). *Bull. Volcanol.* **77**, 34 (2015).

125. Stevenson, R. J., Dingwell, D. B., Webb, S. L. & Bagdassarov, N. S. The equivalence of enthalpy and shear stress relaxation in rhyolitic obsidian and quantification of the liquid-glass transition in volcanic process. *J. Volcanol. Geotherm. Res.* **68**, 297–306 (1995).
<figure><img src="image_7.png" /><figcaption>Extended Data Figure 1 | High- and low-temperature viscosity measurements of samples investigated in this study. See Extended Data Table 2 for experimental data and Extended Data Table 1 for sample chemistry.</figcaption></figure>
<figure><img src="image_8.png" /></figure>

**Extended Data Figure 2** | Comparison between measured and calculated viscosities of all samples at 850 °C. The solid black line is a 1:1 reference line. Symbols are values predicted by the HZ<sup>11</sup> (red) and GRD<sup>12</sup> (black) models. The red and black dotted lines show linear fits to

the model results. The two models generally agree, with the exception of samples J, C and F, which are described better by the HZ model than by the GRD model.
<figure><img src="image_9.png" /><figcaption>**Extended Data Figure 3** | **Raman spectra of samples F, J and L.** These samples show increasing FeO content. Spectra were acquired before the low-temperature viscosity measurements. The peak at about 670 cm⁻¹ (sample L) indicates the magnetite peak. The FeO content (in wt%) is given in parentheses in the key; see Extended Data Table 1.</figcaption></figure>
<figure><img src="image_10.png" /></figure>

**Extended Data Figure 4** | Magnetic-hysteresis analyses of samples F, J, and L. These samples show increasing FeO content. (See Extended Data Table 1 for sample chemistry.) **a**, Hysteresis loops of six sub-samples of samples F and J. *M* is magnetization. **b**, Hysteresis loops of ten sub-samples of sample L. **c**, Hysteresis as in **b** corrected for a paramagnetic

slope calculated using the linear portion of every loop for fields >1.25 T and normalized for the saturation magnetization *M*<sub>s</sub>. The red vertical band indicates all values of B90, calculated for every loop as the field where *M*/*M*<sub>s</sub> = 0.9. The B90 values are used in *d* to estimate the size of the magnetic particles as described in Methods.
<figure><img src="image_11.png" /><figcaption>**Extended Data Figure 5 | Differential thermal analysis data for samples F, J and L.**</figcaption></figure>

These samples show increasing FeO content at a heating or cooling rate of 10 °C min⁻¹ under an argon atmosphere. (FeO content in wt% is given in parentheses in the key; see Extended Data Table 1.) The temperatures given are *T*<sub>g</sub><sup>peak</sup>. Thin lines represent the first heating scans; thick lines represent the second heating scan. The nanolite-bearing sample

(L) shows two endothermic peaks during the first heating scan, revealing two distinct amorphous domains (nanolite-free and nanolite-bearing). The lower glass transition temperature (peak at 692.4 °C) represents the nanolite-free amorphous domain. Except for the topmost curve, the curves were shifted along the y axis for clarity.
<figure><img src="image_12.png" /></figure>

**Extended Data Figure 6** | Total alkali versus silica diagram of all effusive and explosive samples plotted in Fig. 4. Note that the classic total alkali versus silica classification of these samples does not allow us to distinguish between explosive and effusive behaviour. Only through the newly developed RAI can these populations be clearly separated (Fig. 4).
<figure><img src="image_13.png" /><figcaption>**Extended Data Figure 7** | Calculated viscosities at the measured $T_g^{peak}$.</figcaption></figure>

Data reported in Extended Data Table 1 after ref. 125 (see Methods for further details). The viscosity remains constant over the investigated range of silica content with the exception of sample L. This sample exhibits an

extremely high viscosity owing to the iron depletion effect (see main text). The error in the viscosity measurements was estimated to be ±0.05 logarithmic units (see Methods).
<table><caption>Extended Data Table 1 | Chemical composition and iron oxidation state of all glasses (in wt%)</caption><thead><tr><th>Sample</th><th>SiO<sub>2</sub></th><th>TiO<sub>2</sub></th><th>Al<sub>2</sub>O<sub>3</sub></th><th>FeO(T)</th><th>MnO</th><th>MgO</th><th>CaO</th><th>Na<sub>2</sub>O</th><th>K<sub>2</sub>O</th><th>P<sub>2</sub>O<sub>5</sub></th><th>FeO<sub>tilt.</sub><sup>a</sup></th><th>K<sup>#b</sup></th><th>RAI<sup>b</sup></th><th>NBO/T<sup>c</sup></th><th>T<sub>g</sub><sup>peak</sup> (°C)<sup>d</sup></th></tr></thead><tbody><tr><td>A</td><td>77.63</td><td>0.11</td><td>12.73</td><td>3.03</td><td>0.03</td><td>0.06</td><td>0.92</td><td>4.44</td><td>1.62</td><td>0.00</td><td>1.35</td><td>0.19</td><td>0.90</td><td>0.00</td><td>781.0</td></tr><tr><td>B</td><td>77.28</td><td>0.14</td><td>13.39</td><td>2.94</td><td>0.02</td><td>0.06</td><td>0.75</td><td>2.71</td><td>3.61</td><td>0.03</td><td>1.66</td><td>0.47</td><td>0.79</td><td>0.00</td><td>810.4</td></tr><tr><td>C</td><td>77.56</td><td>0.11</td><td>12.42</td><td>1.52</td><td>0.04</td><td>0.08</td><td>0.85</td><td>3.07</td><td>4.08</td><td>0.04</td><td>0.82</td><td>0.47</td><td>0.92</td><td>0.00</td><td>837.8</td></tr><tr><td>D</td><td>76.83</td><td>0.11</td><td>12.43</td><td>2.96</td><td>0.05</td><td>0.07</td><td>0.90</td><td>2.93</td><td>4.29</td><td>0.04</td><td>1.28</td><td>0.49</td><td>0.95</td><td>0.00</td><td>822.4</td></tr><tr><td>E</td><td>75.33</td><td>0.12</td><td>13.61</td><td>2.93</td><td>0.03</td><td>0.07</td><td>0.88</td><td>1.41</td><td>6.80</td><td>0.06</td><td>1.26</td><td>0.76</td><td>0.88</td><td>0.00</td><td>856.6</td></tr><tr><td>F</td><td>78.87</td><td>0.10</td><td>12.52</td><td>1.55</td><td>0.04</td><td>0.06</td><td>0.84</td><td>1.01</td><td>5.28</td><td>0.02</td><td>0.87</td><td>0.77</td><td>0.75</td><td>0.00</td><td>879.8</td></tr><tr><td>G</td><td>77.86</td><td>0.12</td><td>11.69</td><td>2.99</td><td>0.04</td><td>0.07</td><td>0.85</td><td>1.05</td><td>5.34</td><td>0.00</td><td>1.53</td><td>0.77</td><td>0.84</td><td>0.00</td><td>823.8</td></tr><tr><td>H</td><td>77.25</td><td>0.10</td><td>11.95</td><td>2.62</td><td>0.06</td><td>0.22</td><td>1.08</td><td>3.23</td><td>4.35</td><td>0.05</td><td>1.10</td><td>0.47</td><td>1.07</td><td>0.02</td><td>789.4</td></tr><tr><td>I</td><td>76.24</td><td>0.06</td><td>11.46</td><td>2.85</td><td>0.03</td><td>0.44</td><td>1.31</td><td>2.99</td><td>3.83</td><td>0.04</td><td>1.19</td><td>0.46</td><td>1.12</td><td>0.02</td><td>759.1</td></tr><tr><td>J</td><td>73.75</td><td>0.31</td><td>11.99</td><td>3.31</td><td>0.07</td><td>1.64</td><td>2.98</td><td>3.16</td><td>3.56</td><td>0.00</td><td>1.32</td><td>0.43</td><td>1.53</td><td>0.10</td><td>757.6</td></tr><tr><td>L</td><td>71.22</td><td>0.28</td><td>10.86</td><td>6.24</td><td>0.05</td><td>1.58</td><td>2.72</td><td>2.89</td><td>3.28</td><td>0.05</td><td>2.42</td><td>0.43</td><td>1.54</td><td>0.11</td><td>729.8</td></tr></tbody></table>

<footnote>aObtained by wet chemistry titration.</footnote>

<footnote>bSee equation (1) and Methods.</footnote>

<footnote>cNon-bridging oxygens per tetrahedra ratio.</footnote>

<footnote>dMeasured $T_g^{peak}$ with heating or cooling rate of 10 °C min⁻¹.</footnote>
<table><caption>Extended Data Table 2 | Viscosity measurements of all samples</caption><thead><tr><th>Sample name<br>T (°C)</th><th>A</th><th>B</th><th>C</th><th>D</th><th>E</th><th>F<br>log η (Pa s)</th><th>G</th><th>H</th><th>I</th><th>J</th><th>L</th></tr></thead><tbody><tr><td>1591</td><td>3.17</td><td>3.24</td><td>3.51</td><td>3.30</td><td>3.44</td><td>3.70</td><td>3.46</td><td>3.18</td><td>3.05</td><td>2.48</td><td>2.11</td></tr><tr><td>1566</td><td>3.30</td><td>3.39</td><td>3.65</td><td>3.43</td><td>3.58</td><td>3.84</td><td>3.59</td><td>3.32</td><td>3.18</td><td>2.59</td><td>2.21</td></tr><tr><td>1542</td><td>3.44</td><td>3.52</td><td>3.79</td><td>3.56</td><td>3.72</td><td>3.99</td><td>3.73</td><td>3.45</td><td>3.31</td><td>2.72</td><td>2.32</td></tr><tr><td>1517</td><td>3.58</td><td>3.67</td><td>3.94</td><td>3.71</td><td>3.87</td><td>4.15</td><td>3.88</td><td>3.59</td><td>3.45</td><td>2.84</td><td>2.42</td></tr><tr><td>1493</td><td>3.72</td><td>3.82</td><td>4.10</td><td>3.85</td><td>4.02</td><td>4.31</td><td>4.03</td><td>3.74</td><td>3.59</td><td>2.97</td><td>2.54</td></tr><tr><td>1468</td><td>3.87</td><td>3.97</td><td>4.26</td><td>4.00</td><td>4.18</td><td>4.48</td><td>4.18</td><td>3.88</td><td>3.73</td><td>3.10</td><td>2.68</td></tr><tr><td>1444</td><td>4.02</td><td>4.13</td><td>4.42</td><td>4.16</td><td>4.34</td><td>4.66</td><td>4.35</td><td>4.04</td><td>3.88</td><td>3.24</td><td>2.80</td></tr><tr><td>1419</td><td>4.18</td><td>4.29</td><td>4.60</td><td>4.32</td><td>4.50</td><td>4.84</td><td>4.51</td><td>4.20</td><td>4.03</td><td>3.38</td><td>2.93</td></tr><tr><td>1395</td><td>4.34</td><td>4.46</td><td></td><td></td><td>4.48</td><td>4.67</td><td>4.69</td><td>4.36</td><td>4.19</td><td>3.53</td><td>3.07</td></tr><tr><td>750</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>11.20</td><td></td></tr><tr><td>770</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>10.71</td><td></td></tr><tr><td>790</td><td>11.56</td><td></td><td></td><td></td><td></td><td></td><td></td><td>10.98</td><td></td><td>10.30</td><td></td></tr><tr><td>810</td><td>11.10</td><td></td><td></td><td></td><td></td><td></td><td></td><td>10.67</td><td></td><td>9.83</td><td></td></tr><tr><td>830</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td>9.45</td><td></td></tr><tr><td>840</td><td></td><td>11.15</td><td></td><td></td><td>11.52</td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>850</td><td>10.25</td><td></td><td>11.02</td><td></td><td></td><td>11.45</td><td></td><td></td><td>9.90</td><td>9.08</td><td>11.02</td></tr><tr><td>860</td><td></td><td>10.90</td><td></td><td>10.92</td><td>11.25</td><td></td><td>11.58</td><td>10.36</td><td></td><td></td><td></td></tr><tr><td>870</td><td></td><td></td><td></td><td></td><td></td><td>11.18</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>880</td><td></td><td>10.57</td><td></td><td>10.80</td><td></td><td></td><td></td><td>10.13</td><td></td><td></td><td></td></tr><tr><td>890</td><td></td><td></td><td>10.35</td><td></td><td></td><td>10.81</td><td></td><td>9.96</td><td></td><td></td><td></td></tr><tr><td>900</td><td></td><td>10.35</td><td></td><td>10.51</td><td>10.46</td><td></td><td>10.76</td><td></td><td></td><td></td><td>9.91</td></tr><tr><td>920</td><td></td><td>10.02</td><td></td><td></td><td>10.32</td><td></td><td>10.45</td><td>9.33</td><td></td><td></td><td>9.57</td></tr><tr><td>940</td><td></td><td>9.74</td><td>9.61</td><td>9.69</td><td>9.96</td><td>10.05</td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>960</td><td></td><td></td><td></td><td></td><td></td><td></td><td>9.85</td><td></td><td></td><td></td><td></td></tr><tr><td>960</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr><tr><td>980</td><td></td><td>9.10</td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td><td></td></tr></tbody></table>