Our experiments were designed to simulate CO 2 fluxing in magma at mid- to upper-crustal pressure conditions whereby solid carbonate (mainly CaCO 3 but also MgCa(CO 3 ) 2 in some cases) was allowed to react with pre-fused, powdered magmatic rock from Mt. Merapi (Indonesia) and Mt. Vesuvius (Italy) at 1200 °C and 0.5 GPa for up to 300 s (refs 24 and 25). These volcanic systems were chosen because both are subduction-related and display evidence for crustal carbonate assimilation in form of erupted calc-silicate xenoliths and chemical signatures in erupted rocks and fumarole gas22,23,26,27. The advantage of our experiments is that they simulate short-term disequilibrium reactions in order to capture the temporal evolution of magma-carbonate interaction as time variable “snapshots”. Major element compositions of starting materials and experimental products are provided in Supplementary Tables S1 and S2. Boron data for the fused starting materials (n = 25) and the experimental products (n = 147) are provided in Supplementary Tables S3 and S4.

Our experimental products comprise a CaO-normal glass similar in composition to the starting materials, a CaO-rich or MgO-rich glass and a mixing interface between the two domains that shows variable CaO and MgO contents24,25. Incongruent break-down of carbonate produced free CO 2 bubbles that permeated all melt domains28 (Fig. 2). The δ11B values of the starting materials range from −8.8 to −3.5‰ for Merapi and from −14.6 to −7.6‰ for Vesuvius (Fig. 1a). In the Vesuvius case, the measured δ11B values overlap with the lower end of the established Vesuvius range (−7.6 to −6.3‰; ref. 29) and are similar to literature δ11B values for other Italian magmatic systems (e.g., −13.7‰ at Stromboli; ref. 30). The boron concentration of the Vesuvius starting glass ranges from 12 to 14 μg/g and is hence close to the reported range of 14 to 36 μg/g for Vesuvius erupted products29. To the best of our knowledge, there are currently no published δ11B data for Merapi. Boron concentration of our Merapi starting glass ranges from 15 to 18 μg/g, consistent with recorded Merapi whole-rock values of 12 to 20 μg/g (ref. 31), but slightly lower than reported concentrations for Merapi clinopyroxene-hosted melt inclusions (35 to 109 μg/g; ref. 32).

Figure 2 Experimental data. Back scattered electron (BSE) images and δ11B profiles for representative (a,b) Vesuvius and (c,d) Merapi experiments. The solid red line on the BSE images (a,c) represents the SIMS traverse and the red symbols indicate analysis spots. In (b) and (d), average δ11B values for different glass compositional domains are represented by red horizontal bars and the full range of starting material values measured is indicated by black dashed horizontal lines. The CaO-rich glasses have significantly lower δ11B values than both the starting material and the CaO-normal glass due to transport of 11B away from the reaction site in the CO 2 vapour. Error as in Fig. 1. Full size image

Boron isotope profiles were analysed across the interface between CaO-rich and CaO-normal glass (Fig. 2). The δ11B values and B concentration of the CaO-normal glasses range from −5.3 to + 1.2‰ and from 10 to 15 μg/g for Merapi and from −14.7 to −4.9‰ and 9 to 241 μg/g for Vesuvius. In contrast, the δ11B values of the CaO-rich and (MgO)CaO-rich glasses range from −21.9 to −8.6‰ for Merapi and from −41.5 to −13.6‰ for Vesuvius. These values fall considerably below the δ11B values of many subduction systems globally and they are significantly lower than the experimental starting materials (Fig. 1d,e; see also Supplementary Table S5 for sources of all literature data presented). The δ11B values of the CaO-rich glasses are also considerably lower than those of carbonate in the literature. In general, biogenic carbonate has an average δ11B value of +19.1‰ and variable B concentration (Fig. 1). In contrast, lithified carbonate (i.e. lime/dolostone) has lower but still positive δ11B values, which range from +1.5 to +8.4‰ and B concentration of 2 to 18 μg/g (refs 33 and 34; Fig. 1, data sources in Supplementary Table S5). We therefore conclude that simple binary mixing between the reactants cannot explain the low δ11B melts in the experiments and implies additional processes at work.

Boron concentration in the CaO-rich and CaO(MgO)-rich glasses ranges from 4 to 7 μg/g and 2 to 85 μg/g for Merapi and Vesuvius, respectively. A negative correlation between δ11B value and B concentration is observed in all experiments (Fig. 1d,e) and hence the low δ11B values recorded must be related to B degassing from the melt(s) under the experimental conditions (see Methods). We also note that the experimental data (n = 172) for the most part overlap the range of natural magmatic values (Fig. 1), supporting the view that the experiments provide a useful analogue for natural magmatic processes. We hence argue that boron isotope fractionation during decarbonation in the upper plate represents another means to generate low δ11B values in arc magmas in addition to slab dehydration during subduction (cf. refs 3, 8, 9, 10 and 22).

As our experiments were heated to 1200 °C over a period of six minutes (see Methods), the fractionation factor between tetrahedrally and trigonally coordinated boron, α, varied with experiment duration, since α is inversely proportional to temperature13. Realising that α was not constant in our experiments precludes calculation of a unique Rayleigh model for our experimental data. To explain our lowest δ11B values, α probably varied from ca. 1.012 to 1.002 (the latter at 1200 °C), but these values have large uncertainties due to the fact that CO 2 was continually fluxing in the experiments and that fractionation commenced already below 1200 °C. Nonetheless, a similar range of α values may apply to situations where a large thermal gradient exists, such as across metamorphic aureoles in plutonic complexes, or along down-going slabs in subduction zones. Boron isotope fractionation in the experimental melts would also have been mirrored by evolving δ11B values in the co-existing fluid, with δ11B values becoming lower over time in a similar fashion to the slab dehydration models performed by ref. 15. Assuming that boron isotope fractionation is independent of pressure35, our data are in-line with these and similar models that predict values lower than −30‰ in dehydrated subducted materials and others as low as −35‰ for phengite-free dehydrated assemblages at subduction depths15. The fundamental implication here is that extremely low δ11B values can be generated in subducted material at ~3 GPa as well as in magma in the upper arc crust at ~0.5 GPa due to the presence of a coexisting fluid phase that serves to scavenge boron from the rock or silicate melt.

Our experiments also reveal several dynamic aspects of boron transport in magma. In particular, the relatively elevated δ11B values of some of the CaO-normal glass domains compared to the δ11B values of the starting material (Fig. 2) leads us to propose a conceptual model (Fig. 3). In our model, carbonate dissolution and degassing at the onset of magma-carbonate interaction is the catalyst for boron isotope fractionation. The newly formed CO 2 phase scavenges 11B from the carbonate and silicate melt, causing 11B(OH) 3 to enter the newly generated volatile phase by substitution for CO 2 and assimilation of 10B-rich material to occur at the decarbonation reaction site. The highly mobile volatile phase then rapidly migrates away from the reaction site, resulting in the generation of a fluid with a relatively high δ11B value that progressively evolves towards lower δ11B values, similar to some arc fluids and models thereof, as discussed above (Fig. 1c). Conversely, the relatively unaffected, CaO-normal melt further away from the carbonate dissolution site would be undersaturated with respect to CO 2 , which could facilitate coupled 11B and CO 2 reabsorption in the melt and hence relatively high δ11B melt values (Fig. 3). The extent to which this process is expressed as “low” or “high” δ11B melts in individual volcanic systems depends on several factors, including the amount of carbonate assimilated, the viscosity of the magma and, particularly, the solubility of CO 2 in the melt, since low melt solubility of CO 2 will promote bubble formation and thus boron extraction from the co-existing melt. This process would be most effective under low pressures, since CO 2 solubility in silicate melts decreases with decreasing pressure (see discussion in refs 24, 25, 28 and references therein), which would make boron extraction into a CO 2 -bearing phase most efficient in the upper parts of the crust.

Figure 3 Conceptual model. Tetrahedrally coordinated boron is present in carbonate and silicate melt and decarbonation at the onset of assimilation triggers boron isotope fractionation as follows: CaCO 3 (BOH) 4 (carbonate) +SiO 2 (BOH) 4 (silicate melt) → CaO-rich silicate melt +10B(OH) 4 (in melt) + CO 2 (fluid) +11B(OH) 3 (in CO 2 -rich fluid). In other words, assimilation of carbonate into the melt gives rise to Ca-rich melt and a co-existing CO 2 phase that mingles with CaO-normal melt. Transport of trigonally coordinated 11B in CO 2 bubbles away from the reaction site and subsequent partial reabsorption in CO 2 -undersaturated melt at the distal parts of the capsule gives rise to relatively high δ11B values in portions of the CaO-normal glass. Full size image

In conclusion, our data demonstrate that short time-scale (syn-eruptive) carbonate assimilation can result in heterogeneous and locally very low δ11B melt values, similar to predictions for subducted materials. Distinguishing between these processes may be aided by, e.g., the presence or absence of crustal xenolith suites and by thermobarometry to constrain crystallisation depth of the main mineral phases. Boron isotope fractionation in magma via crustal carbonate dissolution is likely to be most pertinent for volcanoes sited on volatile-bearing sedimentary crust, as for instance found in continental subduction settings and may help to identify upper crustal additions to the carbon cycle.