Journal of Volcanology and Geothermal Research 304 (2015) 160–179

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Journal of Volcanology and Geothermal Research

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Review

Volcanoes and the environment: Lessons for understanding Earth's pastand future from studies of present-day volcanic emissions

Tamsin A. Mather ⁎Department of Earth Sciences, University of Oxford, South Parks Road, Oxford OX1 3AN, UK

⁎ Tel.: +44 01865 282125.E-mail address: Tamsin.Mather@earth.ox.ac.uk.

http://dx.doi.org/10.1016/j.jvolgeores.2015.08.0160377-0273/© 2015 The Author. Published by Elsevier B.V

a b s t r a c t

a r t i c l e i n f o

Article history:Received 22 September 2014Accepted 14 August 2015Available online 22 August 2015

Keywords:Volcanic environmental impactsVolcanic gasesTephraSubduction zonesHalogensTrace metals

Volcanism has affected the environment of our planet over a broad range of spatial (local to global) and temporal(b1 yr to 100sMyr) scales andwill continue to do so. Aswell as examining the Earth's geological record and usingcomputermodelling to understand these effects,much of our knowledge of these processes comes from studyingvolcanism on the present-day planet. Understanding the full spectrum of possible routes and mechanisms bywhich volcanism can affect the environment is key to developing a realistic appreciation of possible past and po-tential future volcanic impact scenarios. This review paper seeks to give a synoptic overview of these potentialmechanisms, focussing on those that we can seek to understand over human timescales by studying currentvolcanic activity. These effects are wide ranging from well-documented planetary-scale impacts (e.g., coolingby stratospheric aerosol veils) to more subtle or localised processes like ash fertilisation of ocean biota and im-pacts on cloud properties, atmospheric oxidant levels and terrestrial ecosystems. There is still much to be gainedby studying present-day volcanic emissions. This review highlights the need for further work in three exampleareas. Firstly, to understand regional and arc-scale volcanic emissions, especially cyclingof elements through sub-duction zones, more volatile measurements are needed to contribute to a fundamental and systematic under-standing of these processes throughout geological time. Secondly, there is still uncertainty surroundingwhether stratospheric ozone depletion following volcanic eruptions results solely from activation of anthropo-genic halogen species. We should be poised to study future eruptions into the stratosphere with regard totheir impacts and halogen load and work to improve our models and understanding of the relevant underlyingprocesses within the Earth and the atmosphere. Thirdly, we lack a systematic understanding of trace metalvolatility from magmas, which is of importance in terms of understanding their geochemical cycling and use astracers in environmental archives and of igneous processes on Earth andmore broadly on silicate planetary bodies.Measurements of volcanic rock suites and metals in volcanic plumes have an important part to play in moving to-wards this goal.

© 2015 The Author. Published by Elsevier B.V. This is an open access article under the CC BY license(http://creativecommons.org/licenses/by/4.0/).

Contents

1. Introduction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1612. Studies of the present-day environmental effects of volcanism: a brief synoptic overview . . . . . . . . . . . . . . . . . . . . . . . . . . . 1613. Understanding emission budgets: global, regional and arc-scale processes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 163

3.1. Uncertainties in global volcanic fluxes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1633.2. Understanding regional/arc-scale budgets . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 166

4. The potential of future eruptions to deplete stratospheric ozone . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1684.1. Stratospheric injections of volcanic HCl? . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1684.2. Recent advances in our understanding of volcanic bromine emissions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 170

5. The release of metals from volcanoes: towards a systematic understanding . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 1716. Summary . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 173Acknowledgements . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 175References . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 175

. This is an open access article under the CC BY license (http://creativecommons.org/licenses/by/4.0/).

161T.A. Mather / Journal of Volcanology and Geothermal Research 304 (2015) 160–179

1. Introduction

Volcanism is a key planetary process transferring heat and matterbetween the Earth's interior and its surface reservoirs. The balancebetween these fluxes and other planetary scale processes has been im-plicated in the development of Earth's atmosphere and in maintaininglong-term stability of the surface environment and thus habitability(e.g., Kasting and Catling, 2003). Changes in both short- and long-termrates of planetary outgassing and/or its composition have been connectedwith a range of perturbations to the global environment (e.g., Table 1).These range from the short-lived effects of individual eruptions(Robock, 2000), to the more prolonged proposed consequences of clus-ters or periods of enhanced volcanism occurring both by chance (Milleret al., 2012) and due to silicic ‘flare ups’ (de Silva and Gosnold, 2007) orlarge igneous province (LIP) volcanism (Self et al., 2014a). Other mecha-nisms linking volcanism to global change include more subtle feedbacksbetween the Earth's surface and rates of long-term volcanism or themeancomposition of the gases emitted, for example, in response todegla-ciation or resulting from the subduction of carbonate-rich sediments(Table 1).

Given the rarity or long-term nature of many of these effects, whichboth maintain and perturb Earth's surface environment, much of ourevidence for their characteristics and effects must come from proxyrecords. For example, much can be learnt by combining volatile emis-sions estimates gleaned from melt inclusions in rock samples scaledbydeposit volume/mass (Scaillet et al., 2003)with environmental proxiesin archives such as tree rings (e.g., Briffa et al., 1998; D'Arrigo et al., 2013),ice cores (e.g., Zielinski et al., 1994; Bay et al., 2004; Lavigne et al., 2013;Sun et al., 2014) or sedimentary sections (e.g., Sanei et al., 2012;Percival et al., 2015). Considerable advances have also been made viamodelling studies both of individual past eruptions (e.g., Timmrecket al., 2010; English et al., 2013) and extended events such asfissure erup-tions (e.g., Stevenson et al., 2003; Highwood and Stevenson, 2003; Omanet al., 2006; Schmidt et al., 2010) or LIPs (e.g., Caldeira and Rampino,1990; Beerling et al., 2007; Iacono-Marziano et al., 2012; Black et al.,2013; summarised in Self et al. (2014a)). Comparing similar classes ofpast events that have varying degrees of global environmental impactsis also a powerful approach (e.g., Ganino and Arndt, 2009; Jones et al.,2015). However ifwe are to develop aprocess understanding of planetaryoutgassing and its effects on the Earth's environment it is important toconsider what we can learn from present-day measurements of volcanicemissions and their impacts. It is also timely to consider how futureadvances in our analytical and measurement capabilities might openthe way for a broader understanding in this area.

This review seeks to give a brief overview of what we have learnt todate about the potential effects of volcanism on our planet's surface en-vironment over its geological past by studying present-day volcanism(Section 2). This section is intended to be synoptic and to act as aguide towards the relevant literature to facilitate deeper study of themany research areas covered. I then highlight 3 example areas wherethe scene is set to make substantial progress over the coming years.Section 3 covers the area of global, regional and arc-scale volatile bud-gets. Section 4 discusses the impacts of volcanism on stratosphericozone levels and the challenges of understanding the consequences ofpast volcanism given the uncertainties associated with volcanic plumeprocesses and present-day anthropogenically perturbed stratosphericchemistry. Section 5 examines our understanding of the behaviour oftrace metals in volcanic systems and considers how we might furthersystematise this understanding in the future.

2. Studies of the present-day environmental effects of volcanism: abrief synoptic overview

Present-day active volcanism has a range of chemical and physicalimpacts on the local to regional and global surface environment and at-mosphere of our planet. These effects are largely mediated by the

transfer of significant quantities of different materials from the Earth'sinterior to its surface reservoirs, although volcanic heat has also beensuggested to play a role (e.g.,Mather, 2008). The actual emissions them-selves are studied, in the case of the major volatile gas, aerosol speciesand ash by satellite-based or ground-based remote sensing or in situ di-rectmeasurements. In the case of ash emissions suchmeasurements arecomplemented by mapping and sampling the products left behind onthe Earth's surface by volcanic activity. Melt inclusion and other petro-logical/geochemical studies of these products can also yield key insightsinto volatile loss during volcanism. Current techniques applied to themeasurement of volcanic gas and aerosol are summarised in Table 2.

Similarly a mixture of remote sensing, field studies and experimentsare used to study the environmental consequences of volcanism. Fig. 1and Table 3 summarise the different environmental effects of volcanismsuggested by present-day observations of volcanic events and continu-ous activity. Some of these effects, such as global cooling resulting fromstratospheric sulphate aerosol veils following large explosive eruptionsinto the stratosphere, are relatively well documented. Others, such ashfall to marine and terrestrial ecosystems, the effects of volcanic aerosolon cloud properties and levels of atmospheric oxidants, are more subtleand less well understood. On geological timescales, other mechanismshave also been suggested. These include: (i) significant and prolongedvolcanic CO2 emissions associated with LIP volcanism leading toglobal warming (Chenet et al., 2008; Sobolev et al., 2011; Joachimskiet al., 2012) and ocean acidification (e.g., Payne and Clapham, 2012);(ii) long-term increased uptake of atmospheric CO2 due to increasedweathering resulting from fresh volcanic rock on the Earth's surfacefollowing extended periods of enhanced volcanism (e.g., Schalleret al., 2012; Dessert et al., 2001); (iii) continental ash blanketsincreasing planetary albedo (Jones et al., 2007). The primary associ-ation of these effects with very large-scale volcanism and their spa-tial and temporal scales can make testing their likely extentchallenging using present-day field measurements. Nonetheless suchobservations have a clear role to play. Examples include efforts to im-prove our understanding of C partitioning during igneous processes(see Section 3) and studies of present-day weathering rates at activevolcanoes (Jones et al., 2011).

The limitation of our direct observation of volcanism to the scale thathas occurred in recent times is important to consider when seeking tounderstand the broader interactions of volcanic activitywith the environ-ment. Several studies have shown that ‘scaling up’ from present-day vol-canism to larger-scale events or episodes is not straightforward withfactors such as aerosol microphysics and oxidation agent depletion caus-ing complications (e.g., Bekki et al., 1996;Omanet al., 2006; Schmidt et al.,2010; Timmreck et al., 2010; English et al., 2013). Feedbacks in terms ofenvironmental and climatic impacts via changes in the levels of atmo-spheric oxidants induced by volcanic emissions (e.g., Iacono-Marzianoet al., 2012) are also hard to explore, although there were suggestions ofsuch effects in atmospheric records following the eruption of Pinatuboin 1991 (Manning et al., 2005; Telford et al., 2010). These andother uncer-tainties regarding the environmental effects of volcanism have recentlybeen discussed with respect to LIPs (Self et al., 2014a).

The scientific efforts following the Pinatubo eruption in 1991 (thelargest eruption to the stratosphere to date in the satellite era) yieldedextensive new insights into numerous volcanic impacts on the environ-ment. Historical records, such as those associated with the Laki fissureeruption in 1783 (the closest recent analogue to flood basalt volcanism)pose numerous further hypotheses about such impacts. With unprece-dented arrays of satellite-borne and ground-based sensor networks thescientific community is now better positioned than ever to use currentand future large-scale volcanism to test hypotheses about the responseof climate, atmospheric dynamics and chemistry and feedbacks(e.g., via the response of vegetation, ocean fertilisation or oxidantlevels). The recent response to the 6-month fissure eruption that tookplace at the Holuhraun lava field in Iceland between August 2014 andFebruary 2015 is a good example (Gíslason et al., 2015; Schmidt et al.,

Table 1Examples of modes of volcanism postulated to have the potential to lead to prolonged global change. For further detailed discussion see Mather and Pyle (2015).

Mode Scale of volcanic input Timescale of changed volcanic input Comments

Short-lived large scale explosive eruptions

(e.g., Pinatubo 1991; Tambora 1815;the Young Toba Tuff ~74 ka)

10s–1000s Tg SO2 released.Large-scale stratospheric injection. (McCormick et al., 1995;Oppenheimer, 2003; Chesner and Luhr, 2010)

Hours–weeks. • Clear evidence for global effects following these types of eruption (e.g., Robock, 2000;Oppenheimer, 2002)• Recent modelling studies suggest global temperatures recover on decadal time scales,without tipping the Earth system into a new steady state (Jones et al., 2005; Robocket al., 2009; Timmreck et al., 2010)

Temporal clustering of explosive eruptions

Clusters of relatively small-scaleindividual explosive eruptions

Silicic ‘flare ups’

4 large sulphur-rich explosive eruptions each with globalsulphate loading N60 Tg within 50 years.

Large-scale silicic ‘flare-ups’, e.g., the Altiplano-Puna complex,central Andes: total volume of erupted silicic magma N104 km3

(Mason et al., 2004; de Silva and Gosnold, 2007)

Decadally-paced explosive volcanism overa ~50 year period.

Altiplano-Puna (10–1 Ma)erupted at time-averaged rate~10−3–10−2 km3/yr (de Silva and Gosnold, 2007).

• Sea-ice/ocean feedbacks suggested to lead to the little ice age in ~AD 1350–1850 (Milleret al., 2012)

• Iron fertilisation of the ocean by great volumes of silicic volcanic ash suggested as an effectiveclimatic forcing mechanism that helped establish Cenozoic icehouse conditions coincident withthe ignimbrite flare-up of southwestern North America (Cather et al., 2009)

Large Igneous Provinces (LIPs)

Every 1 km3 of basaltic magma emplaced during a subaerial LIPeruption releases ~3.5–6.5 Mt SO2. Estimates for entire LIPprovinces of 3.5–68 × 106 Tg SO2 (Self et al., 2014a)

Emplacement over several Ma. Detailed studiessuggest most provinces have a ~1 Ma period ofpeak activity with pulses of 10–100 yearsseparated by longer hiatuses (Self et al., 2014a,b)

• Age correlation between at least five LIP provinces of the past 300 Ma with massextinction events and other environmental changes (Wignall, 2001; Courtillot and Renne,2003; Kelley, 2007).• Thought to be due to eruptive gas release from the magmas but exact mechanisms remainelusive• Recent work has highlighted that mantle volatiles might not be the only important factor,e.g., enhanced volatiles from the recycled oceanic crustal component of the mantle plume(Sobolev et al., 2011) and release of volatiles from sedimentary rocks by magmatic heating(Ganino and Arndt, 2009; Svensen et al., 2009).• Fraction of released gases injected into the stratosphere also still debated (Self et al., 2014b)

Feedbacks from the Earth's surface changing the long-term rate of volcanism1

Deglaciation Proposed 2 to 6-fold enhancement in global eruption rates afterthe last glacial maximum (LGM) (Huybers and Langmuir, 2009;Watt et al., 2013a).

5–6 kyr after the LGM (Huybers and Langmuir,2009; Watt et al., 2013a)

• Proposed that associated increases in volcanic emissions may have contributedsignificantly to the post-glacial increase in atmospheric CO2 (Huybers and Langmuir,2009)•Modelling studies of the oceanic carbon cycle (e.g., Roth and Joos, 2012), find only a small rolefor volcanic forcing in post-glacial CO2

• Amore detailed analysis of volcanic eruption records (Watt et al., 2013a) shows sparserevidence for a strong post-glacial volcanic ‘pulse’ and highlights more arc data needed.

Variations in total paleosubductionzone length

Proposed that since the Triassic (250–200 Ma) the totalpaleosubduction-zone length reached up to approx. twicethe present-day value at times (Van Der Meer et al., 2014)

Modelling work suggests variations on timescales≫ 5 Myr (variation wavelengths ~100 s Myr) butpoorly constrained (Van Der Meer et al., 2014)

• Global subduction zone length through time since Triassic quantified using seismictomographic images of deep mantle relics of subduction zones.• Results show agreement with ocean production rates, Sr isotope curves andreconstructed atmospheric CO2 levels.• Increased resolution in tomographic models, more detailed plate reconstructions andimproved constraints on subducting slab carbon content needed to better constrainthe evolution of subduction zones and their individual volcanic degassing outputs(Van Der Meer et al., 2014).

Perturbations to magma volatile load and the composition of the volcanic gases emitted2

Subduction of carbonate-richsediments

Proposed that closure of the Tethys led to up to a ~2-foldincrease in global CO2 emissions due to consequent arc volcanism(Edmond and Huh, 2003; Kent and Muttoni, 2008; Johnston et al.,2011)

~100 Ma during the closure of the Tethys • A poorly constrained portion of subducted C is released to mantle wedge by progressivemetamorphism during subduction.• The paleogeography of the closure of the Tethys ocean suggests subduction of largevolumes of marine carbonate.• The recycling of the C to the atmosphere via volcanism has been postulated as a major factorcontributing to late Cretaceous to early Cenozoic elevated atmospheric CO2 concentrations(Edmond and Huh, 2003; Kent and Muttoni, 2008; Johnston et al., 2011).• Further work is needed to understand the factors controlling the efficiency of C recyclingthrough subduction zones however (Section 3.1; Gorman et al., 2006).

1 Other proposed links between surface forcings such as sea level and the rates of volcanism and its feedbacks through the C cycle remain to be thoroughly tested (e.g., McGuire et al., 1997; Lund and Asimow, 2011; Burley and Katz, 2015).2 Other links between volcanic gas composition and factors such asmantle oxidation state, volatile cycling and average degassing pressure (e.g., Holland, 2002; 2009; Gaillard et al., 2011) have been linked to the oxygenation of Earth's atmosphere.

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2015). There is alsomuch that can be learnt from studies of persistent vol-canic emissions or smaller scale eruptions. In the following sections I pickout 3 examples of broad-reaching scientific questions that offer good op-portunities for progress.

3. Understanding emission budgets: global, regional and arc-scaleprocesses

3.1. Uncertainties in global volcanic fluxes

There are many challenges when determining budgets for volcanicemissions to the atmosphere. SO2 is generally the best quantified volcanicgas species. This is due to its low concentration in the background atmo-sphere, and the availability of straightforward spectroscopic measure-ment techniques that facilitate the measurement of its flux. With somenotable exceptions, where scanning DOAS networks are installed(e.g., Etna, Stromboli, Vulcano, Soufrière Hills Volcano, Tungurahua), theemissions from the majority of volcanoes worldwide are not measuredon a daily basis for prolonged periods even for SO2. Instead they are mea-sured sporadically during field campaigns or periods of heightened activ-ity.When compiling thesemeasurements to yield global or regional-scaleflux estimates the challenge is to combine data that is non-uniformly dis-persed in both space and time and contains sampling biases to producethe best possible assessments. Efforts have been made to compensatefor under-sampling. These have been mainly based on an assumedpower law distribution of volcanic fluxes on a global scale (proposedby Brantley and Koepenick (1995), and used in, e.g., Andres andKasgnoc (1998), and Hilton et al. (2002)). However a recent studyof volcanic emission fluxes in Japan from 1975 to 2006 demonstratedthat the SO2 flux distribution in the cumulative frequency flux plotdoes not obey a power law distribution. This suggests that correctingthe global flux using the power law assumption may not representbest practice. The Japanese dataset was more complete than mostand crucially included measurements from small emitters as wellas the major sources. The cumulative log frequency-log flux plot forthe Japanese dataset was represented not as a straight line but by aroll-off curve bending most significantly at ~500 t d−1 (Mori et al.,2013).

Despite these challenges measurements using ground-basedtechniques and satellite sensors have gone a long way over recentdecades to determining the global SO2 flux. Recent estimates of totalannual emissions range from ~13 to 18 Tg y−1 SO2 (Stoiber et al., 1987;Andres and Kasgnoc, 1998; Halmer et al., 2002) and show reasonableconvergence. For Japan, Mori et al. (2013) report flux data from over 20volcanoes from 1975 to 2006. However, more than 94% of the total fluxwas contributed by 6 volcanoes for most of the period: Tokachi, Asama,Aso, Sakurajima, Satsuma–Iwojima, and Suwanosejima. In other words,the highest flux emitters (greater than a few hundred t d−1) dominatedthe total flux. This highlights that measurement strategies focussed onthe larger continuous or sporadic fluxes are likely to give good global orregional scale SO2 budgets. Similarly, a few volcanoes (including Masaya,San Cristóbal, Pacaya and Fuego) dominate the SO2 flux estimates for theCentral American Volcanic Arc for both the periods 1972–1997 (Andresand Kasgnoc, 1998) and 1997–2003 (Mather et al., 2006). These regionalexamples when extrapolated to the global scale in part explain the conver-gence between estimates of global SO2 fluxes. Essentially if the world'smajor emitters are well constrained then reasonable agreement can be ob-tained despite the incompleteness of the data.

Nonetheless the uncertainties surrounding this global volcanicSO2 flux could certainly be improved and there are major challengesin terms of getting good arc-scale and regional constraints (seeSection 3.1). This is especially true in poorly monitored areas(e.g., Indonesia) which nonetheless form important inputs to regionaland global scale models aimed to understand the impacts of volcanicemissions (Pfeffer et al., 2006; Schmidt et al., 2012). Pushing satellitesensors to measure volcanic SO2 deeper in our atmosphere will be a

crucial component of efforts to improve these inventories. Measure-ments of volcanic SO2 in explosive eruption clouds have been increas-ingly frequent since the 1980s but current sensors lack sufficientsensitivity to reliably capture low-level tropospheric degassing. Recentstudies using the Ozone Monitoring Instrument (OMI) explore the un-certainties surrounding its application to low level, passive volcanicemissions and suggest that it measures only 20–40% of these emissionscompared to ground-based methods (e.g., Carn et al., 2013; McCormicket al., 2013). Investigations are underway to assess the capabilities ofother current instruments such as IASI for such measurements also(Walker et al., 2012). Future instruments such as the Dutch-designedTropospheric Ozone Monitoring Instrument (TROPOMI), scheduled forlaunch aboard ESA's Sentinel-5 Precursor in 2016 have the potential tooffer significant new capabilities in this area. TROPOMI has been de-signed as a replacement for OMI and SCIAMACHY and will offer boththe global daily coverage of the former with the large spectral range ofthe latter. TROPOMI will measure in the UV–vis–near-infrared–short-wavelength infrared ranges, with high spectral resolution and betternadir spatial resolution than either earlier satellite. It is expected toshow even greater sensitivity to SO2 and the focus on tropospheric mea-surements (mainly owing to the mission focus on air quality) is poten-tially advantageous in terms of the measurement of low-level butpersistent volcano emissions into the lower atmosphere.

Although SO2 fluxes are most easily constrained, the fluxes of anumber of other gas species emitted by volcanoes are also of interest.For example, constraining global volcanic carbon fluxes are importantfor a number of reasons. The short- tomedium-term impacts of increas-ing CO2 in surface reservoirs have been much discussed over recentyears and include global warming and ocean acidification. Understand-ing the natural C cycle is important to better understand how humanactivities are perturbing it and informs the debate about the drivers ofcurrent and future global change (Plimer, 2009; Gerlach, 2011). Under-standing this cycle on the present-day planet and in the recent past isalso the key to understanding the coupling between the deep Earthand the surface environment over longer geological timescales(Table 1). Asmentioned above, this coupling can take the form of atmo-spheric maintenance (e.g., Kasting and Catling, 2003) or perturbation(Section 2; Table 1). In other words it has been postulated that volcaniccarbon emissions have played a role in both stabilising Earth's surfaceenvironment and in periods of abrupt environmental change.

Measuring present-day C fluxes in volcanic gases (dominated byCO2) is more challenging than SO2 for a number of reasons. These in-clude the spectroscopic properties of the CO2 molecule, the high levelsin the background atmosphere and the lower solubility of CO2 inmagmas. This lower solubility means that CO2 is exsolved deep withinvolcanic systems and is often lost via diffuse degassing through theground or interactions with hydrological systems in volcanic areas.This diffuse degassing can be as well as or instead of through a focussedplume or fumaroles. Capturing these more spatially extensive diffusefluxes is generallymore challenging than focussed emissions emanatingfrom visible vents or fumaroles.

Vent or fumarole (sometimes termed ‘plume’) CO2 emissions areoften estimated by scaling the SO2fluxwithmeasured C/S ratios. The re-cent advent of MultiGas and similar systems has greatly extended therange of systems where C/S plume measurements have been deter-mined (Table 2). These C/S ratios are then also used to scale up to globalplume CO2 fluxes using the global volcanic SO2 flux. Thus uncertaintiesin the global volcanic SO2 flux propagate to global CO2 flux estimates.However, additional uncertainties are introduced due to the choice ofrepresentative C/S ratios (see Section 3.2) and CO2 flux emissions,such as diffuse soil degassing, that are decoupled from SO2 emissions.Measurement of diffuse CO2 emission fluxes are usually achieved viasoil accumulation chamber measurements (e.g., Chiodini et al., 1998)although other methods such as tunable lasers are more infrequentlyapplied (Pedone et al., 2014). Localised or grid measurements are thenscaled up to the total area of the volcanic structure using systematic

Table 2Examples of methods for measuring volatile gases and aerosol emissions from present-day volcanism.

Species/parametermeasured

Techniques Comments Example references

Gas fluxes (SO2) Ground-based traversing or scanning byCOSPEC/DOAS spectroscopic methods for loweraltitude plumesSatellite U.V. or I.R. sensors for larger/highervolcanic plumes

Due to its low background levels and spectroscopic properties thevolcanic gas flux of SO2 is most commonly measured. Permanentnetworks can give high frequency measurements over prolongedperiods.

Ground-based: Reviewed inMcGonigle and Oppenheimer(2003).Galle et al. (2010)Satellite: Reviewed in Carn et al.(2013) and McCormick et al.(2013). Carboni et al. (2012)

Major gas speciesratios with SO2

⁎Bulk plumesH2O/CO2: FTIR or Multi-GasHCl/HF: FTIR or alkali traps in filter packs orother device followed by lab analysis (e.g., ionchromatography).

FumarolesGiggenbach flasks for all species as well as bulkplume methods

Used to scale up to fluxes of other major gas species from those ofSO2. Trace gases like HBr and HI can be measured in alkali trapsusing more sensitive techniques like ICP-MS.Multi-Gas or similar sensors can be ground-based or mounted onUAVs or manned aircraft.

New more recent innovations are discussed in Wittmer et al.(2014). Robustness is a strong consideration when working involcanic environments.

FTIR: Burton et al. (2007)Multigas: Shinohara et al. (2008)Filter packs: Aiuppa et al. (2005a)Giggenback flasks: Taran et al.(1995)

Diffuse soildegassing (CO2)

The accumulation chamber method is mostcommonly used.

There are challenges surrounding scaling up from a discontinuousmeasurement grid to an appropriate flux. Areas of degassing mayalso be inaccessible or too large to cover densely.

Chiodini et al. (1998)

Particle numbersand sizedistribution

Optical particle counters (in situ)Sun photometers and other remote sensingmethods.

Reviewed in Table 2 of Matheret al. (2003)

Particle chemistry Direct sampling of volcanic plume particles ontofilter papers or traps including the use ofimpactors to gain size resolved information.

Reviewed in Table 2 of Matheret al. (2003)

⁎ Ifmeasurements of volcanic emissions are not feasible then petrological studymelt inclusions or volatile-bearingminerals (e.g., apatite) in the volcanic rockproducts can oftenbeusedto gain insight into the melt parental volatile contents and degassing.

Fig. 1. Schematic summary of volcanic processes observed to have environmental impacts from studies of present-day volcanism. Two example volcanic scenarios are presented, namelyash-free passive degassing to the troposphere and substantial explosive volcanism with an ashy Plinian column punching into the stratosphere. In reality volcanic activity can have char-acteristics of both these scenarios and its nature can vary through an eruptionwith time.Major gas species in volcanic plumes at the point of emission are H2O, C (mainly as CO2), S (as SO2

andH2S depending on P, T and redox conditions) and halogen halides (examples of these are given in the explosive plume). The numbers in red circles refer to different volcanic processesthat have environmental consequences. Their consequences are expanded upon in Table 3 with reference to these numbers. Other long-term/larger-scale effects of volcanism(e.g., prolonged CO2 emissions, enhanced weathering of fresh volcanic material and increased albedo due to ash blankets) have also been postulated and are discussed in Section 2 ofthe text.

164 T.A. Mather / Journal of Volcanology and Geothermal Research 304 (2015) 160–179

Table 3A summaryof the environmental effects of volcanism forwhichwe have foundevidence through recent studies of present-day volcanism. Other long-termeffects have also been postulated(see Section 2 of the text for discussion).

Volcanic process(numbers refer to those shown in Fig. 1)

Measured environmental effects Example references

Explosive volcanism

(1) Oxidation of volcanic SO2 to sulphate aerosol inthe stratosphere

• Sulphuric acid aerosol in the stratosphere interacts withincoming solar radiation increasing planetary albedo.• After the 1991 Pinatubo eruption average cooling of thetroposphere lasted 1–3 years.• Evidence that this cooling impacts global precipitation rates.• Interaction of the aerosol with incoming and outgoingradiation warms the stratosphere leading to dynamicresponses of the atmosphere such as northern hemispherewinter warming.• Associated increase of diffuse (as opposed to direct) radiationsuggested to impact forest photosynthesis rates althoughevidence is not conclusive.• Change in stratospheric UV flux, temperature and thepresence of aerosol surfaces have potential to perturbstratospheric chemistry (see 2).

Reviewed in Robock (2000). Krakauer and Randerson(2003); Trenberth and Dai (2007)

(2) Heterogeneous chemistry on/in particles in thestratosphere

• Presence of surfaces for heterogeneous reactions affects O3

reaction cycles.• Quantifying the global scale effects of volcanic eruptions onstratospheric ozone challenging - chemical and dynamic effectsoccur simultaneously and effects are not much larger thannatural variability.• Depletions may be dependent on the presence ofanthropogenic chlorine in the stratosphere (see discussion inSection 4).• Understanding of the importance of reactions on ash surfacesespecially in the high-T regime of dense explosive eruptionplumes is in its infancy.

Reviewed in Robock (2000); Ayris et al. (2013; 2014)

(3) Ash fall onto land • Near to eruption ash has a clear physical effect blanketingecosystems, choking drainage etc. and can effect therespiratory health of animals following inhalation.• Ash releases various chemical species to the environmentfollowing deposition and contact with rain or surface water.• Recent results from Chile suggest explosive volcanism influ-ence on structure and composition of temperate rainforestsover the last 10,000 years.

Witham et al. (2005); Horwell and Baxter (2006); Stewartet al. (2006); Martin et al. (2009); Henríquez et al. (2015)

(4) Ash fall into sea • Volcanic ash fall into the Earth's oceans brings with it awide range of chemical elements that can have impacts onthe marine environment including macro- andmicro-nutrients.• The delivery of volcanic ash to Fe-limited waters has beenhypothesized as a natural fertilisation mechanism forphytoplankton.• Increased marine productivity has been proposed to leadto enhanced drawdown of atmospheric CO2.

Reviewed in Duggen et al. (2010) and Browning et al.(2015). Hamme et al. (2010); Langmann et al. (2010);Hoffmann et al. (2012); Olgun et al. (2013); Achterberget al. (2013); Browning et al. (2014); Mélançon et al.(2014)

(5) Aerosol settling from the stratosphere • Evidence for the possible effect of stratospheric volcanicaerosols on seeding cirrus cloud formation exists forindividual cases following tropopause folds.• Initial studies of the global significance of such effectsfollowing Pinatubo suggest not a significant climatefeedback on the global scale.

Reviewed in Robock (2000). Sassen et al., (1995); Luo et al.(2002); Campbell et al. (2012)

(6) Volcanic lightning and particle charging • Electrostatic charging likely exerts control on fall outpatterns of ash.• Plume lightning will produce reactive species such as NOx

that may then take part in further reactions.

Reviewed in Mather and Harrison (2006). Rose et al.,(2006); Martin and Ilyinskaya (2011)

Passive degassing

(7) Oxidation of volcanic SO2 to sulphate aerosol inthe troposphere

• Sulphate aerosol is present in volcanic plumes from thepoint of emission and further sulphate is generated duringadvection of the plume downwind following theentrainment of atmospheric oxidants.• This aerosol can have detrimental effects on vegetation andhuman/animal health if it fumigates surface environments,provides surfaces for heterogeneous chemistry (see (9)) andinteracts solar radiation directly as well as effecting cloud coverand properties (see (11)).

Reviewed in Mather et al. (2003; 2004). Sutton and Elias(1993); Schmidt et al. (2011)

(continued on next page)

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Table 3 (continued)

Volcanic process(numbers refer to those shown in Fig. 1)

Measured environmental effects Example references

(8) Plume gas fumigation and particle settling toground-level

• Leaf damage observed downwind of passively degassingvolcanoes due to fumigation by acid gases.• Chemical and magnetic evidence for particle and gasuptake by leaves.

Reviewed in Delmelle (2003). Quayle et al. (2010); Martinet al. (2012).

(9) Heterogeneous chemistry on/in particles in thetroposphere

• Particular recent progress understandinghalogen-mediated destruction of ozone and formation ofBrO.• Involves heterogenous reactions in/on particles (acidaerosol), the entrainment of atmospheric oxidants into theplume and photochemistry.• Thought to be initiated by trace reactive chemical speciesfrom high-T reactions in the volcanic vent or duringlightning flashes in more explosive eruptions.

Reviewed in von Glasow et al. (2009) and von Glasow(2010)

(10) Rain through the plume • Rain falling through volcanic plumes strips out soluble gasspecies, aerosol and any ash present.• Leads to acid rain (with similar detrimentalenvironmental effects to anthropogenic acid rain) and therelease of other chemical species such as trace metals to theenvironment.

Delmelle (2003); Calabrese et al. (2011); Mather et al.(2012)

(11) Mixing with the background atmosphere • Many possible reactions as a volcanic plume mixes withthe background atmosphere or anthropogenic pollution.• Recent evidence of modification of cloud occurrence andproperties by volcanic aerosol with important implications forplanetary radiative budget.

Gassó (2008); Yuan et al. (2011); Schmidt et al. (2012);Ebmeier et al. (2014)

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assumptions about how fluxes vary spatially (e.g., variationswith volca-nic activity, fault locations, observed ground alteration; Chiodini et al.,1998; Parks et al., 2013; Hutchison et al., 2015). The balance betweendiffuse and plume emissions appears to vary depending on the eruptivestate of the volcano. Hernández et al. (2015) showed that volcanoeswith a recurrent eruptive period shorter than 10 years generally showa diffuse/plume ratio in the range 0.1–1,whereas thosewith a longer re-current eruptive period (~100 years) show an exponential increase ofthis ratio. The state-of-the-art regarding present-day volcanic C fluxeswas recently reviewed by Burton et al. (2013). They compiled previousestimates of global volcanic subaerial CO2 flux as between 66 and300 Mt/yr and revised this flux upwards to 540 Mt/yr by includingdiffuse flank emissions (117 Mt/yr), emissions from volcanic lakes (94Mt/yr) and emissions from tectonic, hydrothermal and inactive volcanicareas (66 Mt/yr) in addition to the more usually quantified plume emis-sions (271 Mt/yr). More recent estimates have suggested that theplume emission estimate for volcanic CO2 in Burton et al. (2013) maybe too high as it depends on what method you use to correct for unmea-sured fluxes (Pedone et al., 2014). This highlights the need for furthermeasurements of both plume and diffuse fluxes (see also Section 3.2)and improved understanding regarding how best to scale up spatiallydispersedfluxes such as diffuse volcanic CO2 emissions. Satellitemeasure-ments might one day also play a role for volcanic CO2 measurements al-though there are extensive challenges to overcome first (e.g., Heymannet al., 2015). Large, high-altitude explosive plumes will likely offer themost promising initial targets rather than lower flux, surface level diffusedegassing.

Recent reviews also exist regarding the fluxes of volcanic halogens(e.g., Pyle and Mather, 2009), which are of interest in terms environ-mental impacts such as acid rain and ozone depletion (Fig. 1). In allcases data frommore volcanic systems as well as systematising our un-derstanding of the degassing process is crucial in terms of improvingthese estimates (see also Section 4).

3.2. Understanding regional/arc-scale budgets

While there is some convergence, at least for SO2, in the global volca-nic flux estimates, to further our understanding of several of themodes of

volcanic impact detailed in Table 1we requirefluxes onmore specific tec-tonic or geographical scales. For example, the balance between the fluxesfrom arc volcanoes compared to those in other tectonic regimes is impor-tant to quantify given evidence of different responses to deglaciation be-tween tectonic settings (Huybers and Langmuir, 2009;Watt et al., 2013a).Similarly, a key issue in the understanding of the cycling of volatiles be-tween the surface and the mantle is volatile processing through subduc-tion zones (Luth, 2014; Schmidt and Poli, 2014). One method employedto understand the processes involved in this cycling and its efficiency isby comparing estimated input volatile fluxes associated with thesubducted slab with estimated output degassing fluxes (e.g., Hiltonet al., 2002; Fischer, 2008; Johnston et al., 2011).

As discussed above because of the relative ease of its measurementSO2 fluxes are generally the best constrained for global volcanoes. Thisis also true on smaller geographical scales. Several studies have compiledSO2 fluxes fromdifferent arcs and then used these fluxes and X/SO2 ratios(where X is another element or compound present in volcanic emissions)to scale up to the arc-scale fluxes of other chemical species such as CO2

(e.g., Hilton et al., 2002; Mather et al., 2006; Fischer, 2008; Taran, 2009;Johnston et al., 2011). Complementary techniques use melt inclusionsand petrological modelling of the geochemistry of the eruptive productsof arc magmatism to calculate output fluxes (e.g., Wallace, 2005;Freundt et al., 2014). Both approaches have their issues but in generalagree that higher fractions of the volcanic arc outputs of H2O and Cl arecontributed by the downgoing subducting slab than for CO2 (Straub andLayne, 2003; Wallace, 2005; Mather et al., 2006; Taran, 2009; Freundtet al., 2014). Similar estimates for S are less common, in part, as generallythe input flux on the downgoing slab is not as well constrained. Thewidespread in the subduction zone recycling efficiency estimates for S (i.e., theratio of inputs to outputs) reflects sensitivity to this poorly constrainedinput (Taran, 2009; Freundt et al., 2014).

Hilton et al. (2002) made one of the first attempts to compiledegassing data on an arc-by-arc basis to probe the comparative volatilerecycling efficiencies through the different global arcs. Johnston et al.(2011) then used these data to explore the relationship of CO2 recyclingefficiency to the thermal structure of the subduction zones in question,and argued that their results supported the proposal that changes in Ce-nozoic climate could be related to the closure and subduction of the

Fig. 2. Different estimates of the SO2 flux from different global subduction zones. Subduc-tion zones are distinguished according to their thermal parameter (Syracuse et al., 2010).Where this parameter varied along the arc a mean value was taken. Arc lengths are takenfrom those compiled inHilton et al. (2002) other than the exceptions described in the text.The filled red diamonds are the estimates used in Hilton et al. (2002) and Johnston et al.(2011) fromAndres and Kasgnoc (1998). The open black squares are themedian of the es-timates compiled in Halmer et al. (2002) with the error bars representing the ranges.These estimates include explosive and passive emissions. Thefilled blue square representsthe mean estimate for the CAVA from Mather et al. (2006) with error bars indicating therange. Filled blue circles represent the estimates for Japan from Mori et al. (2013) withthe higher value including the period of intense degassing by Miyakejima and the lowervalue not. The filled blue triangle is that value for Kamchatka–Kuriles estimated byTaran (2009). All values are without correction for unmeasured fluxes.

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Tethys Ocean (Edmond and Huh, 2003; Kent andMuttoni, 2008). In lightof these studies it is useful to review the current state-of-the-art in termsof the global arc-scale degassing budgets made from present-day mea-surements. This will allow us to assess the validity of this approach andwhat work we need to prioritise in order to promote progress in thefuture.

I will discuss arc-scale SO2 fluxes first. Hilton et al. (2002) extrapo-lated SO2 fluxes by filtering the degassing fluxes compiled in AndresandKasgnoc (1998) for each arc. They then used a power law correctionproposed by Brantley and Koepenick (1995) to account for unmeasuredSO2 fluxes. Hilton et al. (2002) used the continuously erupting fluxeslisted in Andres and Kasgnoc (1998) augmented in some cases byother sporadic fluxes. The challenges associated with combining fluctu-ating continuousfluxeswith large sporadic short-livedfluxes to achievemeaningful volcano SO2 flux estimates on various temporal and spatialscales are discussed above in the global context. In terms of arc-scalemeasurements lessons can again be learnt from the well-monitoredareas of the world. As mentioned above, Mori et al. (2013) recentlycompiled all available SO2 flux data from 32 years (1975–2006) ofmon-itoring at Japanese (Japan, Nankai, Izu-Bonin and Ryukyu) volcanoes(93 volcanoes corresponding to about 10% of the world's arc volcanoesby number). The activity of these volcanoes waswell recorded over thisperiod. Using these data they were able to evaluate the temporal varia-tion of the flux of each volcano in some detail. This analysis unveiled themagnitude of variability from year to year possible over this length ofarc (3950 km or ~9% of total global trench length). Their maximum es-timated annual arc-scale fluxwas 9.2 Tg yr−1 in 2001, with a minimumof 0.8 Tg yr−1 in 1981. This variability of over an order ofmagnitudewaslargely due to the intense degassing of Miyakejima volcano after 2000.The average annual flux for the period is 2.2 Tg yr−1 with Miyakejimaincluded and 1.4 Tg yr−1 with Miyakejima emissions omitted. Thus asingle high emitter is capable of significantly skewing regional time-averaged degassing totals.

Fig. 2 compares the different estimates for arc-scale fluxes for Japanand globally. In order to aid inter-arc comparisons the totalfluxes are nor-malised to arc length although this does introduce a further uncertainty interms of the appropriate arc length to choose. Arc lengths were takenfrom those compiled in Hilton et al. (2002). They were similar tothose used in Johnston et al. (2011) other than (i) for Japan where theNankai, Izu-Bonin and Ryukyu arc lengths were all included here and(ii) Indonesiawhere Java, Sumatra and E. Sundawere summed for our es-timates. The length for the Andes only includes Colombia and Peru(2550km). Amore realistic value (N7000 km) includes Chile, and reducesthe fluxes per km by about a third. Fig. 2 shows that based on currentmeasurements, we can constrain the time-averaged SO2 fluxes fromwell-studied arcs (e.g., Japan) to within a factor of ~2. However, for lesswell-studied arcs this is not the case. Hilton et al. (2002) used theAndres and Kasgnoc (1998) compilation of data from only 3 volcanoesto estimate a flux of 0.1 Tg a−1 from the 1900 km long Kamchatka–Kurilearc. Amore recent and complete survey (Taran, 2009)with data from N20volcanoes suggests that the arc-scale flux is N1.1 Tg a−1, more than anorder of magnitude higher. Estimates are also very sensitive to themode of degassing included in compilations. For the Indonesian arc,which hosts N70 active volcanoes, Halmer et al. (2002) estimated anaverage annual SO2 flux for Indonesia of 2.1–2.6 Tg yr−1 including bothexplosive and quiescent activity. This is again over an order of magnitudegreater than arc fluxes based solely on passive emissions (0.1 Tg yr−1:Hilton et al., 2002). A similar situation is observed for Alaska (Fig. 2).Quiescent and continuous emissions from both these arcs are alsounder-sampled due to inaccessibility (e.g., Bani et al., 2015).

Thus, while on the global scale there is some convergence in our un-derstanding of time-averaged volcanic SO2 emissions, caution is neededwhen interpreting arc-scale emissions budgets. Not only is more workneeded to understand how best to include short-lived periods of highflux in time-averaged arc-scale fluxes, but we are also dealing withdata that is sparse in time and space and biased towards accessible

volcanoes. Satellite measurements offer an attractive solution to theseproblems given their global coverage. As discussed in Section 3.1 thereare still limitations to the application of satellite volcanic SO2 in thelower atmosphere but better understanding of the associated uncer-tainties and future sensors like TROPOMI provide exciting newopportu-nities in this area. More extensive continuous volcanic SO2 monitoringat more volcanoes, better compilation and archiving of data and im-proved methods for dealing with under-sampling will also allow us tomake meaningful progress in terms of constraining arc-scale SO2 fluxesover the coming decades.

As discussed above fluxes of other chemical components of volcanicdegassing are often scaled to those of SO2. Therefore, as for global esti-mates, uncertainties in arc-scale SO2 fluxes (i.e., a minimum of a factorof 2 even in well studied arcs) propagate to such flux estimates forother plume components. Again, when scaling up to arc-scale emissionsof other plume components a further complication comes into play,namely choosing appropriate X/SO2 ratios. Measurements of theseratios again tend to be temporally and spatial sparse and focussed onwell-monitored systems. Taking C/S ratios for example, studies atwell-monitored volcanoes show that carbon and sulphur degassingis poorly correlated in many cases, e.g., due to deep exsolution anddiffuse rather than plume emission of CO2 (discussed in Section 3.1).At Etna, one of the best monitored volcanoes worldwide and a majorvolcanic CO2 source, the C/S ratio of the gases varies over short time-scales (~months) by a factor of N2 and over longer timescales by a factorof N~30 (Aiuppa et al., 2007; Shinohara et al., 2008). At Mt St Helensduring the 1980s a ~10-fold variation in the C/S was noted (Harrisand Rose, 1996). As well as fractional degassing, processes such as vari-able scrubbing of SO2 by hydrothermal systems and sulphide breakdownin magmas are sources of variability in C/S ratios (e.g., Edmonds et al.,2013). Understanding the degassing processes in operation at each indi-vidual system is important for the interpretation of C/S ratios.

Hilton et al. (2002) used different C/S molar ratios for the differentarcs based on measurements made at each, largely at fumaroles. Thesevalues ranged from 1.7 for Kamchatka–Kuriles to 6.5 for Japan and 8.5for Italy (higher measured values for New Zealand and the Philippinesof 28.7 and 104.5 were not used but were replaced by a ratio of 5).Johnston et al. (2011) used a slightly different range of arc ratios basedon the compilation of fumarole measurements in Fischer (2008). Thesevaried from 0.32 and 0.92 for Alaska–Aleutians and Kamchatka–Kuriles

Fig. 3. The mean and range of molar ratios of C/S in volcanic plumes from different arcs.The arc fumarolemeasurements (filled squares) are taken from the compilation in Fischer(2008) made using the Giggenbach flask method. The arc ‘plume’ measurements (filledcircles) include some plumes that originate from fumaroles (e.g., Vulcano, Soufrière HillsVolcano) but many are from open vents or lava lakes. ‘Plume’ measurements for non-arcvolcanoes (open circle) are also illustrated. ‘Plume’measurements are based on FTIRmea-surements orMultigasmeasurements. The numbers below the x-axis indicate the numberof volcanoes wheremeasurements were taken and that themean and range ratio is basedon. References: CAVA (Masaya: compiled in Martin et al., 2010, Telica and Turrialba:Conde et al., 2014), Andes (Villarrica: Shinohara and Witter, 2005; Sawyer et al., 2011,Lascar: Tamburello et al., 2014, Lastarria: Tamburello et al., 2014), Italy (Stromboli: Burtonet al., 2007, Vulcano: Aiuppa et al., 2005b; McGonigle et al., 2008), Japan (Aso: Mori andNotsu, 1997, Mikakejima: Shinohara et al., 2003), Kamchatka (Gorely: Aiuppa et al.,2012), Antilles (Soufrière Hills Volcano: Edmonds et al., 2010), Mexico (Popocatépetl:Goff et al., 2001), Vanuatu (Yasur: Métrich et al., 2011), non-arc volcanoes (Erebus:Oppenheimer and Kyle, 2008, Erte 'Ale: Sawyer et al., 2008a, Nyiragongo: Sawyer et al.,2008b, Etna: Aiuppa et al., 2006; 2008, Kīlauea: Gerlach et al., 1998 and Edmonds et al.,2013 and references therein). Measurements from explosive eruptions in Alaskasuggested mean mass CO2/SO2 values of 0.8–1.3 with higher values attributed to hydro-thermal scrubbing of SO2 (McGee et al., 2010 and references therein).

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respectively, to 6.60 and 17.87 for the Andes and Italy respectively. Fig. 3compares the Fischer (2008) C/Smeasurementswith thosemadewith al-ternativemethodologies frommain ‘plumes’ rather than individual fuma-roles. The ratios span a great range of values for individual volcanicsystems, within arcs and globally. While there may well be real variationin the C/S ratios between arcs, current data coverage (Fig. 3) is not suffi-cient to distinguish this. For example, the range of fumarole C/S ratiosmeasured in Japan (Fig. 3) is almost as great as the global range. In ad-dition, although the fumarole measurements for the Andes (fromGaleras) suggest a higher C/S ratio than some other arcs might beappropriate for the Andes, this is not borne out by the ‘plume’measure-ments (fromVillarrica, Lascar and Lastarria). In both ‘plume’ and ‘fuma-role’ datasets in Fig. 3 there is some suggestion that Italian C/S fluxes areat the higher end of global compilations. This may be attributed to base-ment rock decarbonation at in the Mediterranean region (e.g., Parkset al., 2013; Tassi et al., 2013). While interesting in itself, this is distinctfrom control on C/S ratios by deeper subduction driven processes. Similarprocesses have been suggested to explain the C degassing systematics atPopocatépetl (Goff et al., 2001) that also plots with a high C/S ratio inFig. 3.

Compiling the arc-by-arc uncertainties in SO2 flux and C/S ratios itbecomes apparent that, while broad-brush global estimation of volcanicCO2 flux is improving, the degassing data on an arc-by-arc scale muststill be treated with caution. These data are not yet fit for purposewhen trying to understand the factors affecting the efficiency of Crecycling by comparing arc to arc CO2 gas emission and much work re-mains to be done before we can claim to have a quantitative under-standing of present-day let alone past subduction recycling scenarios.Large uncertainties are therefore associatedwith estimates of individualarc CO2 recycling efficiencies obtained from gas measurements to date(Johnston et al., 2011). Future work should aim to use current andnew technologies to improve arc-scale degassing measurements of allgas species (e.g., MultiGas/gas sensor boxes, Aiuppa et al., 2014;Tamburello et al., 2014). Understanding subduction zone volatile pro-cessing will require the combination of more extensive gas data withholistic budget and process understanding from multiple approachesincluding experimental techniques, petrological methods, exhumedsubduction zone studies and innovative new modelling (e.g., Connolly,2005; Gorman et al., 2006; Dasgupta and Hirschmann, 2010;Marín-Cerón et al., 2010; Watt et al., 2013b; Shinohara, 2013; Agueand Nicolescu, 2014; Kelemen and Manning, 2015). Although giventhe complexity of the processes thermodynamic and geodynamicmodelling of the behaviour of C through subduction zones is in its infan-cy, initial modelling studies do suggest that both arc-thermal structureand sediment thickness, are indeed important, with sheer input amountof C into the subduction zone also a key player (Gorman et al., 2006;Fig. 4).

4. The potential of future eruptions to deplete stratospheric ozone

The widespread identification of ‘mutated’ forms of fossil spores andpollen in end-Permian rocks has led to recent claims that the SiberianTrap volcanism, coincident with the Permo-Triassic mass extinction(251 Ma), led to a worldwide depletion of stratospheric ozone(Beerling et al., 2007). Modelling studies to explore the possibility ofsuch large-scale ozone depletion suggest the delivery of magmatichalogens to the stratosphere from Siberian Trap volcanism is an impor-tant atmospheric input to consider (Beerling et al., 2007; Black et al.,2013). LIP volcanism presents particular challenges when consideringthe stratospheric effects of associated halogen emissions. The explosivityof trap volcanism like that in Siberian is uncertain and thus estimates ofthe fraction of LIP emissions that would have reached the stratosphereare not reliable (Self et al., 2014b). Further, in addition to the issues ofobtaining good estimates ofmagmatic halogen degassing, the importanceof emissions of organohalogens due to magmatic interactions with coun-try rock and other indirect sources is poorly constrained for these intense

and prolonged volcanic episodes. However, even for explosive plumesthat are known to reach the stratosphere,much remains to be understoodabout the fate of chlorine and other volcanic halogens and hence their po-tential effects on stratospheric ozone. Here again there is much to learnfrom studying present-day volcanism.

4.1. Stratospheric injections of volcanic HCl?

Stratospheric ozone depletion was attributed to the Pinatubo erup-tion in 1991. Estimates of column depletions ranged from about 2% inthe tropics to about 7% in the mid latitudes and reached about 20% inthe aerosol cloud itself. This depletion has been attributed to anthropo-genic chlorine in the stratosphere being activated via heterogeneouschemistry involving volcanic aerosols (summarised in Robock (2000)).No substantial HCl (in general the most abundant halogen in volcanicemissions) increase was detected in the stratosphere following thiseruption (Mankin et al., 1992; Wallace and Livingston, 1992), despitepetrological estimates of HCl emissions of ~4.5 Mt HCl (Westrich andGerlach, 1992). This indicates efficient removal of HCl in the Pinatuboeruption column before it could enter the stratosphere (likely byscavenging of soluble gas into hydrometeors) and suggests that in thepre-anthropogenic era volcanic eruptions would not have causedstratospheric ozone depletion. However, measurements from ElChichón and smaller eruptions since Pinatubo present a more complexpicture (Table 4). Mid latitude and tropical eruptions show evidencefor significant HCl injection into the stratosphere (up to 0.04 Tg for ElChichón) despite estimated scavenging rates of 80–99.8% duringplume transit through the lower atmosphere (Table 4). Measurementsfrom the high latitude eruption of Hekla in 2000 show little evidencefor HCl removal in the lower atmosphere (although the compositionof the source gases is poorly constrained). Localised areas of completeozone destruction were observed within the plume during aircraft en-counters with the Hekla 2000 plume in the stratosphere (Rose et al.,2006). In all cases in Table 4 where HCl mixing ratios are reported,

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maximum in-plume levels (2.5–70 ppbv)werewell above natural back-ground levels and comparable to the anthropogenically perturbedstratosphere.

Modelling studies also show some variation in terms of the predict-ed stratospheric HCl injection from explosive volcanic eruptions.Modelling by Tabazadeh and Turco (1993) suggested that HCl partitionsalmost completely into supercooled drops, leading to a reduction in thegas phase concentration of up to four orders of magnitude. However, allthe hydrometeors in the Tabazadeh and Turco (1993) model wereliquid, and incorporation of gases in ice particles was not included.Textor et al. (2003) used a different plume model and suggested thatthe ice phase was in fact dominant. This reduced liquid scavenging ofHCl in the model and suggested that scavenging by ice particles was asignificant process. Many of these particles reach the stratosphere andthey concluded that, despite scavenging, more than 25% of the HClreaches the stratosphere. The results of both these modelling studiesare consistent with observations to date of mid latitude and tropicaleruptions (Table 4) but not with the high latitude Hekla case. TheTextor et al. (2003) model runs used subtropical atmospheric conditionvalues, so it is perhaps not surprising that there is a better fit to lowerlatitude data. Despite progress in themodelling of HCl behaviour in vol-canic plumes, there are still important processes that have yet to be fullyunderstood, parameterised and included in themodels. For example re-cent experimental studies highlight the complex role of HCl uptake ontovolcanic ash as a possible mode of HCl transport to the stratosphere aswell as a mode of removal and deposition (Ayris et al., 2014).

Fig. 4. (a) and (b) Recycling efficiency (%) at different subduction zone calculated based on them(Syracuse et al., 2010) – note the log scale – and (b) sediment thickness (Plank and Langmuir(d) Recycling efficiency (%) at different subduction zone calculated fromGorman et al. (2006) pldecreasing C recycling efficiency for colder (i.e., higher thermal parameter) slabs. Thick sedimenrecycling efficiency at a given thermal parameter. Conversely in some arcs C is recycled more elow sediment thickness and a relatively highC input to the subduction zone held in the oceanic cper km into the subduction zone is also seen with thick C-poor sediments pulling Andaman anbroad trend is discernible from these plots. Gorman et al. (2006) predicted a global slab-derivewell with the lower end of the estimate range for total global subaerial volcanic CO2 fluxes (se

The strength of volcanic HCl stratospheric injection is critical to in-form our understanding and modelling of the environmental impactsof past, current and future volcanism. For example, Beerling et al.(2007) allowed 75% of the HCl emitted by the Siberian Traps to reachthe stratosphere in their model. They argued that this was appropriatedue to the high latitude of this volcanic activity making it more analo-gous to the Hekla 2000 scenario (Rose et al., 2006) than the modellingstudies of Tabazadeh and Turco (1993) and Textor et al. (2003). Their re-sults would have likely been different if they had used the 0–25% strato-spheric injection rates suggested by these modelling studies and theother observations summarised in Table 4. However without a process-based understanding for the likely stratospheric injection rate under dif-ferent conditions it is hard to identify a most likely scenario. Progress inthis area will rely on improved understanding in several inter-relatedareas of research. Firstly,wemust better understand the initial abundanceand release of Cl frommagmatic systems and any associated shallow con-tributions involved in different volcanic scenarios (e.g., Pyle and Mather,2009; Svensen et al., 2009). New analytical techniques, opening up theuse ofmineral phases such as apatite, offer exciting newpotential in addi-tion to improved melt inclusion methodologies (e.g., Stock et al., 2015).Secondly, furtherwork is needed to understand the processing of chlorinewithin the plume once it is emitted. This will involve further explorationof the sensitivity of stratospheric injection to different environmental(e.g., ambient atmosphere water content, tropopause height, atmo-spheric dynamics) and volcanological factors (e.g., explosivity) usingcurrent models (e.g., Textor et al., 2003; Glaze et al., 2015). It will also

odelling results presented in Gorman et al. (2006) plotted against (a) thermal parameter, 1998) on the downgoing slab. (c) Subducting C flux in downgoing slabs (Jarrard, 2003).otted against total C input per kmof trench (Jarrard, 2003). There is a general trend towardst (e.g. S. Antilles) and high input of C held in sediment (e.g. CAVA, Colombia) tend to reducefficiently than their thermal parameters would suggest. At Tonga this can be explained byrust rather than sediments. A broad trendof reducingC recycling efficiencywith total inputd S. Antilles off this trend. No immediate reasons for the deviation of Kermadec from thed C flux of 0.35–3.12·1012mol/yr (15–140Mt/yr) from convergent margins which agreese text and Burton et al. (2013)).

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require further field measurements and laboratory experiments to un-derstand poorly parameterised plume processes and facilitate theireventual inclusion in the models (e.g., gas uptake by hydrometeorsand ash and their fate: Ayris et al., 2014 and Martin et al., 2012 and ref-erences therein). Measurements of HCl and other halogen species in thestratosphere following futuremajor eruptions are also to be encouraged(see Table 4 and Section 4.2 for discussion of Br measurements).

As an aside, we can use themeasurements presented in Table 4 to es-timate the time-averaged HCl flux to the stratosphere. If we assume anaverage stratospheric volcanic SO2 flux of 1–4 Tg SO2/yr (Pyle et al.,1996) and a mass ratio into the stratosphere of HCl/SO2 of between0.006 and 0.06 (Table 4), this yields a time-averaged HCl flux to thestratosphere of 0.006–0.24 Tg/yr.

4.2. Recent advances in our understanding of volcanic bromine emissions

Chlorine is not the only halogen released by volcanoes with the po-tential to deplete ozone. Bromine is thought to be at least an order ofmagnitude more effective in terms of destroying stratospheric ozonethan chlorine (Daniel et al., 1999). The behaviour of volcanic bromine inmagmas, duringdegassing andonce in the atmosphere ismorepoorly un-derstood than chlorine. This is due in part to the greater analytical chal-lenges associated with its measurement. These are largely associatedwith its lower concentrations in both rocks and volcanic gases (Bureauet al., 2000; Bureau and Métrich, 2003; Pyle and Mather, 2009). Recentanalytical advances allow for optimism that the next decades will see astep change in our understanding of volcanic Br emissions. Recent inno-vative measurements of volcanic bromine include:

(a) Gas phase BrO (a product of ozone destruction by Br)

BrO has been measured in both ‘passive’ or mildly explosive tro-pospheric plumes (Stromboli, Etna, Villarrica, Tungurahua, Masaya,Soufrière Hills Volcano, Nyiragongo, Sakurajima and Ambrym) byground-based remote sensing (Bobrowski et al., 2003; von Glasowet al., 2009 and references therein) and in more ash-rich plumes fromexplosive eruptions (Kasatochi, Eyjafjallajökull and Puyehue) via air-borne or satellite platforms (plumeheights 8–14 km in the upper tropo-sphere/lower stratosphere region: Theys et al., 2009; Heue et al., 2011;Theys et al., 2014). Hörmann et al. (2013) used the satellite-basedGOME-2 instrument to systematically look for BrO during moderate tomajor eruptions and very strong degassing events. Plumes from 10 to12 different volcanoes were found to contain BrO of volcanic origin

Table 4Summary of explosive volcanic plume HCl measurements around the tropopause or in the stra

Volcano Measurementsa Altitude(km)

El ChichónMarch/April 1982(Mankin and Coffey, 1984;Bluth et al., 1997)

HCl: 0.04 TgSO2: 7 TgGives a mass ratio = 6 × 10−3

N11.9 km

HeklaFeb 2000(Rose et al., 2006)

SO2 ~ 1ppmvHCl ~70ppbvMean mass ratio = 0.068 (from paper)

10.4 km a.s.l(lowerstratosphere)

Soufrière Hills VolcanoMay 2006(Prata et al., 2007)

0.003–0.01 Tg HCl0.2 ± 0.02 Tg SO2

Gives mass ratio = 0.014–0.056

Stratosphere

Puyehue-Cordón CaulleJune 2011(Theys et al., 2014)

HCl up to 2.5 ppbvMolar HCl/SO2 = 0.03–0.11 based onmixing ratios reported.Mass ratio = 0.02–0.06

12–14 km a.s.(aroundtropopause)

a Note for comparison that the natural background level of Cl in the stratosphere is about 0.b Pyle andMather (2009) compiled direct plumemeasurements fromhigh-temperature emissio

Erebus was higher (phonolite) but other volcanoes from non-arc settings tended to be lower. Thi

(~30% of all volcanoeswhich emitted SO2 plumes detectable by the sen-sor during this period). These included several examples (Dalaffilla,Kizimen, Kliuchevskoi, Nabro and Sarychev) where BrO was detectedfor the first time.

(b) Gas-phase HBr (believed to be the major gas species emitted bymagma)

This has been quantified in the tropospheric plumes from the arcvolcanoes Masaya, Telica and Villarrica (Witt et al., 2008; Sawyeret al., 2011) and the non-arc volcanoes Etna, Kīlauea and Nyiragongo(Aiuppa et al., 2005a; Mather et al., 2012; Bobrowski et al., 2015) viacollection on filter packs and analysis using ICP-MS.

(c) In situ glass concentrations of Br

The first measurements of Br in melt inclusions and matrix glass bysynchrotron–X-ray fluorescence were recently presented from 14 largeexplosive eruptions over the last 70 ka from Nicaragua (Kutterolf et al.,2013). A further study presenting measurements across the broaderCentral American volcanic arc seeks to explore the arc-scale origins ofvolcanic Br (Kutterolf et al., 2015).

There have been many key lessons from these initial studies. Theysuggest that Br is about 3 orders of magnitude less abundant in volcanicgases than Cl (Pyle and Mather, 2009; Mather et al., 2012; Kutterolfet al., 2013) and that our current best estimate of the average flux tothe atmosphere of HBr is 5–15 Gg/a (Pyle and Mather, 2009). The gas-phase and melt inclusion measurements of Br suggest that it is likelyenriched in subduction-related emissions due to the contribution of thedowngoing slab (Mather et al., 2012; Pyle and Mather, 2009; Kutterolfet al., 2015), although recent measurements from the non-arc volcanoNyiragongo suggest further complexity here (Bobrowski et al., 2015).Themelt inclusion data suggest that large individual eruptions can poten-tially emit at least up to 1100 kt Br (Kutterolf et al., 2015) and the satellitebased measurements show that volcanic Br can be transported to theupper troposphere/lower stratosphere region by even relatively small ex-plosive eruptions (e.g., Theys et al., 2009; Heue et al., 2011; Theys et al.,2014). The measurements of BrO suggest that Br is involved in heteroge-neous chemistry within the plume following emission with associatedimpacts on atmospheric ozone levels (von Glasow et al., 2009).

The scene is set but there ismuch still to be done. As discussed aboveour understanding of Br behaviour in volcanic systems is in its infancy

tosphere since 2003.

Lat./Long. Comments

22°N to35°N

• Using arc mean HCl/SO2 mass ratio of 0.3b suggests scavengingefficiency of ~98%.• Aircraft FTIR measurements using the sun as a source through theupper atmosphere

76.4–75.7°N9.0–4.4°W

• Little evidence of HCl scavenging (HCl/SO2 mass ratios tend to belower for non-arc volcanoes)b

• In-situ HCl measurements made by CIMS on board an aircraft

Tropics • Based on Soufrière Hills Volcano HCl/SO2 ratio of 0.34–6.80b givesscavenging efficiencies of 84–99.8%• HCl measurements made by satellite microwave limb sounder (MLS)

l. 40.59°S72.12°W

• Using arc mean HCl/SO2 mass ratio of 0.3b suggests scavengingefficiency of 80–93%• HCl measurements made by satellite microwave limb sounder (MLS)

6 ppbv and levels peaking at ~3.7 ppbv around 1997 (Nassar et al., 2006).ns ofHCl/SO2. Arc gasesweremass ratio 0.1–0.3with geometric/flux-weightedmean of ~0.3.s compilation did not include measurements of opaque explosive plumes.

171T.A. Mather / Journal of Volcanology and Geothermal Research 304 (2015) 160–179

compared to Cl. Future studies focussing on the content and degassingbehaviour of Br in melts (Bureau et al., 2000; Bureau and Métrich,2003) and building up our database of Br emissions data (especiallythe major species HBr) are to be encouraged. Modelling or laboratorystudies to understand the fate of Br once in the atmosphere would beof great value (Cadoux et al., 2015). Key issues in terms of understand-ing the potential of volcanism to impact stratospheric ozone surroundhow much volcanic Br might reach the upper atmosphere for differentemissions scenarios and under different atmospheric conditions.

5. The release of metals from volcanoes: towards a systematicunderstanding

Volcanic emissions are an important source of trace metals to theenvironment (e.g., Hinkley et al., 1994; Pyle and Mather, 2003). Forsome metals such as Hg and Cd they may account for up to 50% of natu-rally occurring emissions (Nriagu, 1989). Understanding the release ofthese often toxic trace metals from volcanoes followed by their transportin volcanic plumes is important. Many of them impact human health andthe environment (e.g., Fig. 1,Martin et al., 2009) and are important signalsin archives of volcanic activity and environmental impacts such as icecores (e.g., Hinkley andMatsumoto, 2007; Kellerhals et al., 2010). For ex-ample, recentwork onHg has shown that volcanogenic tracemetals havegreat potential in global chemostratigraphy, and might thereby help un-ravel in even finer detail the links between past episodes of large-scalevolcanism and their global environmental consequences (e.g., Saneiet al., 2012; Percival et al., 2015). An understanding of metal volatilityand release from magmas is also important due to its role in the for-mation of some ore deposits (although this is thought to be generallylimited without hydrothermal processes e.g., Pokrovski et al., 2013, andreferences thereinWilliams-Jones andHeinrich, 2005). On a larger scale,tracemetal volatility from silicatemelts is key to understanding theiruse as tracers of igneous/magmatic processes andwider processes asso-ciated with the composition, evolution and differentiation of silicateplanets (e.g., Wood et al., 2006; Edmonds, 2015).

However, the processes by which trace metals are released frommagmas and the factors controlling their release are poorly understood.Once a gas phase is present other minor elements such as trace metalscan diffuse through the liquid and partition into the vapour. Thispartitioning behaviour is likely dependent on temperature, pressure,magmatic composition and gas composition (which in the case ofsurface degassing may include an atmospheric component). Thus, thedegree to which trace metals migrate and partition into the gas phaseis a function of their diffusion rate in the liquid, affinity for the vapourphase and equilibrium solubility. A systematic understanding of metalvolatility is desirable. This section addresses how studies of present-day volcanism might further contribute to this overall aim.

While there has been recent progress in terms of experimental stud-ies on the behaviour of volatile trace metals during degassing frommagmas, much remains to be done. Mackenzie and Canil (2008) exper-imentally measured diffusivities for Te, Tl, Re, Cd, Pb and Sb in a simpleFe-free ‘haplobasalt’melt composition. They proposed that trace metalswith different diffusivity and partition behaviourwill fractionate duringbubble growth and degassing and that metal ratios such as Cd/Re couldbe used as a precursory signal of volcanic eruptions. Johnson and Canil(2011) investigated the degassing of volatile heavy metals (Bi, Pb, Tl,Au, Re, Sb, Sn, Cd, Mo, As, Cu) from doped natural basalt anddacite and synthetic rhyolite melts over a range of temperatures(1200–1430 °C) exposed to air at 0.1 MPa. They showed that the addi-tion of Cl and S increased by two-to five-fold the volatilities of all metals,with S having amore profound effect. Limited spectroscopic plumemea-surements (CuCl(g): Murata, 1960), laboratory experiments (Ammanet al., 1993) and thermodynamic modelling (e.g., Symonds et al., 1992)have previously suggested that metals are generally volatilised fromthe magma as chloride species. Johnson and Canil (2011) also highlight

the importance of diffusion aswell as partitioning andmelt compositionas a control on the trace metal composition of the emitted gases.

The lower concentrations of metals in volcanic plumes continue tobe an analytical challenge and blank levels can be an issue in somecases. This has led to many previous studies of metals in volcanicgases focussing on condensates and sublimates collected a volcanic fu-maroles and thermodynamic modelling (reviewed in Pokrovski et al.(2013)). While valuable, in many cases these fumarolic studies are astep removed from the actual volatilisation of the metals from themagma. Between emission from themagma andmeasurement at the fu-marole, processes such as cooling of the gas stream, gas-wallrock and hy-drothermal interactions etc. might have occurred. Over the past fewdecades more data has been amassing on metal emissions measured di-rect frommagma–air interfaces such as lava lakes, vents and flows. Therehave also been technological developmentswith techniques such as ICP-MS nowbeing routinely available andwe have learntmuch about the re-leased of trace metals from magmas via such measurements at activevolcanoes. With some notable exceptions (e.g., Hg, Pyle and Mather,2003) the majority of trace metals in volcanic plumes are present bythe time of measurement in the particle phase (Hinkley, 1991). Metalsmeasured in volcanic plumes will be from 2 main sources: (i) release/volatilisation of metals from the magma followed by gas-particle phaseconversion and (ii) a silicate phase. The silicate phase will include com-ponents from fragmentation of themagma during explosions or bubblesbursting at its surface and erosive dust from the conduit walls (reviewedin Mather et al., 2003). These 2 pathways from magma in the vent tometals measured on the particle filter have different implications interms of the propensity of a metal to be volatilised and emitted intothe environment independent of ash or lava.

To disentangle the contributions from ash/silicate versus volatilisationand thus assess the volatility of trace metals frommagmas several meth-odologies have been used including: (i) emanation coefficients (εx),(ii) enrichment factors (EFs) and (iii) weight ash fraction (WAF). Thesedifferent methods are summarised in Table 5. Each of them has their ad-vantages and disadvantages.While emanation coefficients should be eas-ily comparable between different systems, their nature means that theycan only be directly measured under rare circumstances (e.g., Rubin,1997; Norman et al., 2004). Even then it is hard to know how to be surethat the final concentration, Cf, has truly been measured. Further,although the orders of element volatilities can be compared, it is notclear how useful emanation coefficients are in terms of comparing withexperimental results. Johnson and Canil (2011), for example found littlecorrelation between their results and emanation coefficients (maybe asthey were measuring diffusion coefficients). Enrichment factors (EFs)are more regularly possible to estimate but can vary considerably evenat the same volcano on the same day due to factors such as variableash/silicate content of plumes. This, and the different choices of referenceelement between studies, are challenges in terms of comparing the vola-tility of different metallic elements from different volcanic systems. Thismakes developing a systematic framework to explain differences intheir behaviour (e.g., as a function of magma composition) challenging.Comparison of the order of element EFs betweendifferent systems shouldyield useful if qualitative information however. Another method that hasbeen used to estimate the ash contribution to trace metals on particlefilters is the weight ash fraction (WAF, Table 5) defined in Aiuppa et al.(2003). This has many advantages, although it relies on good qualitymeasurements of the rare Earth elements (REEs) in both the magmaand plume phases and the (likely relatively safe) assumption that theREEs are not volatilised. In the absence of reliable REE data, the concen-trations of other non-volatile elements likeMg or Al could be used in thecalculation instead. It is also by definition an estimate of the fraction ofash in the sample and is not a measure of metal volatility that can becompared between systems. Aswell as improving the number andqual-ity of measurements of metals in high-temperature volcanic plumes,considering howbest to compare these results quantitatively to developa unified framework of metal volatility is also a key ongoing challenge

Table 5Summary of some of the main methods used in the literature to assess metal volatility in volcanic eruptions.

Method Definition Comments

Emanationcoefficients(εx)

εx ¼ Ci−C fCi

where Ci is the initial concentration of element x in the magma andCf is the final concentration of element x in post-eruptive lava.

• Defined by Gill et al. (1985) and Lambert et al. (1985/86) in order to studythe degassing of radionuclides such as 222Rn, 210Po, 210Bi and 210Pb• Challenging to quantify εx values for elements that have intermediateproperties. By considering disequilibrium with a radioactive decay seriesLambert et al. (1985/86) determined values for Bi and Pb and Pennisi et al.(1988) used these to calculate other εx values1

• This makes the determination of all other emanation coefficients reliant onhow well the value of εPb is determined and care should be taken about whento make an ash/spatter correction in the calculation• Rubin (1997) used this approach to calculate and compile global averageemanation coefficients for a number of elements

Enrichmentfactors(E.F.)

EF ¼ ðXplume=YplumeÞðXmagma=YmagmaÞ

where X is the element of interest and Y is a reference element

• As noted by numerous other authors (e.g., Crowe et al., 1987; Aiuppa et al.,2003), these values will vary even at the same volcano on the same day• One of the main factors leading to this variation is the variable ash/silicatecontent of volcanic plumes even over short timescales2

Weight ashfraction(WAF)

WAF ¼ XerXt

¼ ðREEt Þa� ðX=REEt ÞrXt

where Xer is the erosive contribution to Xt the total amount of element X in theaerosol, (REEt)a is the total rare earth elements measured on the particle filterand (X/REEt)r is the average ratio of element X to the total rare earth elementsin the parent rocks

• Defined by Aiuppa et al. (2003) this allows the subtraction of the ‘noise’ ofash contribution from total element content in order to gain insight into tracemetal degassing mechanisms• Other non-volatile elements could be used in place of the REE

1 By considering disequilibriumwith a radioactive decay series Lambert et al. (1985/86) determined values for Bi and Pb. Pennisi et al. (1988) then used the values for values for Bi and

Pb (Lambert et al., 1985/86) and the following relationship to calculate other εx values:ðX=PbÞgasðX=PbÞlava ¼

εx=εpbð1−εx Þ=ð1−εpbÞ.

2 Ifwe assume, for an open vent system, that the elements in theplumederive from2 routes: (i) volatilisation from themagma and (ii) emission of ash or other silicatematerial from themagma surface then Xplume= Xvol+ Xsil, whereXvol is the component ofX contributed by the volatile route and Xsil is the component contributed from the silicate phase and similarly Yplume

= Yvol+ Ysil. Thus the equation for E.F. becomes:EF ¼ ððXvolþXsilÞ=ðYvolþYsilÞÞ

ðXmagma=YmagmaÞ . If Y is verymuch less volatile than X (e.g., whenMg or Al are taken as reference elements, Mather et al., 2012) then

increasing the amount of ash collected on thefilterwill decrease the EF as the increasing silicate componentwill have a proportionally greater effect on (Yvol+ Ysil) than (Xvol+ Xsil) all elsebeing equal. Conversely if Y is verymuchmore volatile than X (e.g., when Br is taken as the reference element, Crowe et al., 1987) EFwill increase if the amount of ash collected on the filterincreases.

172 T.A. Mather / Journal of Volcanology and Geothermal Research 304 (2015) 160–179

for scientists working in this area. It is this challenge and possible futureways to resolve it that I focus upon for the rest of this section.

Most geochemical processes, including degassing are modelled byuse of partition coefficients. Most of the experimental data on metalmelt-fluid/vapour partitioning is for geological systems at moderatepressures and high temperatures (P = 0.1 to 2 kbar; T = ~1000 °C;see e.g., Pokrovski et al. (2013), and references therein) with littledata existing for low-pressure scenarios relevant in terms of magmaticdegassing. However, the broadly used concept of partition coefficientswas recently applied to field observations of high-temperature volcanicdegassing of trace metals. Zelenski et al. (2014) estimated partitioncoefficients from the effusive basaltic eruption of Tolbachik volcano,Kamchatka by taking the ratio of the metal concentrations in the highconcentration gas stream (corrected for dilution by air) with elementalconcentrations measured in the rock. They made similar estimates forErta Ale and Kudryavy for comparison citing uncertain contributionsfrom the erosive/ash component to the aerosol and reference lavavalues as among possible factors accounting for differences. Thisseems like a promising approach to compare metal volatilities mea-sured at different volcanic systems and with experimental data. In anextension of Zelenski et al. (2014)'s approach, I suggest correcting forthe variable contribution from the silicate fraction in the sample usingthe WAF in order to consider only the volatile-derived component inthe aerosol:

Conc. element X in volatile derived aerosol = (1-WAF) × Total conc.element X in aerosol.

In order to estimate a ‘partition coefficient’ for degassing for X bynormalising each element in the gas phase to its concentration in themagma, these concentrations have to be corrected for dilution prior tomeasurement. Zelenski et al. (2014) used a dilution factor derivedfrom the ratio of total pumped air to condensate to correct back tomag-matic gas concentration upon emission from the magma. Alternativelythe measured mass ratio X/S (i.e., the volatile component of element X

in the sample divided by the concentration of gaseous S ideally mea-sured on alkali impregnated filters within the same filter stack) multi-plied by the concentration of S (in g g−1) in the gas mixture at thepoint of emission can be used. This requires knowledge of the gas com-position for that particular volcano at least in terms of the major gasesH2O, CO2 and SO2 and any other significant contributions. Typical gascompositions can be used if suitable simultaneous measurements areunavailable. In Table 6 I apply thismethodology (aerosolmeasurementscorrected using WAF and dilution before comparison with rock mea-surements) to estimate ‘volatility coefficients’ (i.e., approximatedegassing partition coefficients) for metals in the gas and rock from 2example volcanic systemswhere suitable published data exist: Etna vol-cano, Italy (Aiuppa et al., 2003) andMasaya volcano, Nicaragua (Mouneet al., 2010). Table 7 summarises the results of these calculations andcompares them to those in Zelenski et al. (2014). The normalisation torock composition would ideally be based on a sample from the lavaflow/lake or shallow reservoir at a similar time to the gas samplingbut, as with EF calculations in the past, this is hard to achieve due topractical constraints.

These estimates are by no means without uncertainties. For example,it requires good quality geochemical measurements of a range of volcanicproducts and, aswith EF andWAF above, it is sensitive to the rock compo-sition chosen. This rock composition itself can varywith processes such asmagma source, thermal/crystallisation and degassing history of themagma. Nonetheless this preliminary exploration shows promise as aunifying framework and there are interesting similarities and differ-ences between the results from Etna, Masaya, Tolbackik, Erta Ale andKudryavy (Table 7). If we define volatile elements as those with volatil-ity coefficient N1, moderately volatile with 0.1–1 and non-volatile thosewith b0.1, we see broad agreement between the elements falling intothese categories between the different volcanoes. This is especially thecase if we consider volatile and moderately volatile as one group(Table 7). As, Bi, Cd, Hg, Re, Se, Tl and In all have volatilities N0.1 ineach case where they are measured. Al, Ba, Ca, Co, Cr, Fe, Ga, Ge, Hf, K,Li, Mg, Mn, Na, Nb, Ni, Rb, Sc, Si, Sn, Sr, Th, Ti, U, Y, Zn and Zr are non-

Table 6Volatility coefficients (estimate of concentration in gas/concentration in rock at point ofemission) for Etna and Masaya calculated from published data (Etna: Aiuppa et al.,2003; Masaya: Moune et al., 2010).

Element Volatility coefficient

Etna⁎

Mean (s.d.)+Masaya$

Mean (s.d.)

Ag 3.5 × 10−2 (8.1 × 10−2)Al 1.1 × 10−2 (1.5 × 10−2)As 3.3 × 10−1 (2.8 × 10−1)Au 3.0 × 10−2 (5.1 × 10−2)Ba 0§ 0Be 3.7 × 10−2 (7.5 × 10−3)Bi 3.8 (1.8) 23 (4.2)Ca 1.1 × 10−3 (1.6 × 10−3)Cd 1.6 × 10−1 (7.7 × 10−2) 10.4 (2.1)Co 1.9 × 10−2 (2.3 × 10−2)§ 0Cr 7.2 × 10−3 (6.4 × 10−3)§

Cs 9.0 × 10−2 (1.7 × 10−1)§ 2.0 × 10−1 (3.8 × 10−2)Cu 1.5 × 10−1 (1.7 × 10−1) 1.2 × 10−1 (2.2 × 10−2)Fe 0Ga 9.5 × 10−3 (1.3 × 10−2)Ge 1.0 × 10−3 (9.7 × 10−4)Hf 0§ 0Hg 19#

In 3.0 × 10−1 (1.8 × 10−1)K 1.6 × 10−2 (2.8 × 10−2)Li 9.7 × 10−3 (2.5 × 10−2)§

Mg 6.4 × 10−4 (9.0× 10−4)Mn 3.4 × 10−3 (2.9 × 10−3)Mo 9.8 × 10−3 (5.7 × 10−3)Na 7.4 × 10−3 (1.5 × 10−2)Nb 0§

Ni 6.8 × 10−4 (9.4 × 10−4)§

Pb 1.3 × 10−1 (1.4 × 10−1)§ 5.9 × 10−1 (1.6 × 10−1)Pt 3.8 × 10−1 (4.1 × 10−1)Rb 3.1 × 10−2 (5.6 × 10−2)§ 7.0 × 10−2 (1.5 × 10−2)Re 7.3 (4.9)Sb 17 (30) 1.6 × 10−1 (4.5 × 10−2)Sc 1.5 × 10−3 (1.4 × 10−3)§

Se 9.2 (8.4)Si 0Sn 1.3 × 10−3 (1.8 × 10−3)Sr 0§ 0Te 1.5 × 10−2 (1.2 × 10−2) 26 (7.3)Th 8.5 × 10−6 (1.5 × 10−5)§ 1.4 × 10−3 (8.9 × 10−4)Ti 1.1 × 10−3 (1.4 × 10−3)§

Tl 9.8 (13) 20 (3.3)U 4.6 × 10−3 (6.2 × 10−3)§ 4.7 × 10−2 (1.3 × 10−2)V 4.6 × 10−4 (6.8 × 10−4)§

W 3.2 × 10−3 (6.8 × 10−3)Y 0 0Zn 1.4 × 10−2 (1.2 × 10−2) 4.4 × 10−2 (1.2 × 10−2)Zr 0§ 0

⁎ X/SO2 ratioswere calculated assuming SO42−/SO2 ~0.01 bymass. Etna gas datawas taken

from Aiuppa et al. (2008). Etna is an alkali-basaltic volcano formed on tensional lithosphericfaults cutting the eastern margin of the Apennine–Magrehebide chain in the Ionian Sea.

+ In some cases one or more filters of the set were not estimated to have any volatilecomponent on them or yielded WAF N 1. These were assigned volatility coefficients of 0.

# Taking bulk plume Hg/SO2 ~8.7 × 10−6 by mass from Bagnato et al. (2007) as Hg ispredominantly in the gas rather than particle phase at the point of emission. 4.2 × 10−1

was calculated for the aerosol alone.§ To test the sensitivity of the results to the rock data used I repeated the analaysis using

the trace element composition data for ash erupted from Etna's North East Crater May–December 1995 (Schiavi et al., 2006) for all the elements indicated. Pb was the only elementthat would change categories in Table 7 based on this re-analysis with a new volatility coef-ficient of 9.9×10−2. This is stillwithin the standarddeviationof the results from thedifferentfilters. This comparison is unlikely to capture the effects of using an undegassed magma.

$ X/SO2 ratios were calculated from gas filter data in the Moune et al. (2010) paper.Masaya gas data was taken from Martin et al. (2010). Masaya is a basaltic shield volcano,forming part of the Central American Volcanic Arc associated with the subduction of theCocos plate beneath the Caribbean plate.

173T.A. Mather / Journal of Volcanology and Geothermal Research 304 (2015) 160–179

volatile in all cases measured. For the elements Ag, Au, B, Cs, Cu, Mo, Pb,Pt, Sb, Te andW there is some disagreement in their volatility betweensystems, although some like Pb are only borderline non-volatile in a sin-gle case and Mo is only borderline moderately volatile in a single case.

This might reflect uncertainties in the data. Alternatively these may bereal differences due to magmatic or fluid phase compositions, redoxconditions and temperature for example, and the complexity of the pro-cesses of metal degassing from magmas and fluid transport(e.g., Johnson and Canil, 2011; Vlastélic et al., 2013; Johnson et al.,2013; Henley and Berger, 2013; Edmonds, 2015).

It is interesting to compare the data presented in Table 7 to othersystems for quantifying elemental volatility. If we group the emanationcoefficients detailed in Rubin (1997) into 3 classes with (i) εx N 10−2,(ii) εx 10−4–10−2 and (iii) εx b 10−4, the elements break down intothe categories as follows:

(i) B, Bi, Se, Cd, Re, Hg, Au, Ag, Pb(ii) Cu, Ir, As, Te, Zn, Tl, In, Mo, Sn, W, Sb, Cs, K, Ni, Li, Na(iii) Hf, Cr, Co, Rb, Sr, V, Ce, Fe, Yb,Mn, Ba, Sc, Th, Eu, Ca, La, Sm, Al, Mg.

There is broad agreement with Table 7, although also some interest-ing exceptions. For example, the εx of B suggests that it is more volatilethan the volatility coefficients suggest and the εx of Tl suggests that it isless volatile. For both these elements the Kudryavy volatility coefficientsare more consistent with the εx than those for the other systems. Therelatively high volatility of B suggested at Kudryavy and in Rubin(1997) is also in contrast to the findings from a recent melt inclusionstudy at Kīlauea (Edmonds, 2015). Fig. 5a also highlights that Te andSb show greater volatility in terms of their volatility coefficients com-pared to εx. Based on EFs, Moune et al. (2010) grouped elements in tohighly volatile (Cd, Tl, Bi and Te) and moderately volatile (Be, U, Zn,Rb, Cu, Sb, Cs and Pb). Again there is general consistency, but in our vol-atility coefficient analysis U, Zn and Rb are defined as non-volatile andSb belongs in the most volatile group for Etna. On a broader planetaryscale, the condensation temperatures (Tc, at which 50% of an element'satoms in a protoplanetary gas disc would have condensed) calculatedby Lodders (2003) are often used to probe the role of volatility in thechemical evolution of the Earth and silicate planets. Comparing thesewith our estimate of metal volatilities based on measurements atdegassing volcanoes suggests, for example, that Ge, Rb, Sn and Zn areless volatile under these circumstances than their Tc values suggest,while Pt, Sb and Re are more volatile (Fig. 5b). These differences likelystem from the very different conditions considered in the Lodders(2003) calculations which were for condensation from gas in thelow pressure, reduced conditions in the planetary disc rather thanvolatilisation from present-day magmas into an oxygenated atmosphere.

These differences between volcanoes and volatility classificationsystems highlight the importance of a systematic understanding ofmetal volatility under different conditions when considering their be-haviour in the Earth system in the past, present and future. This alsohighlights the importance of developing a unified framework to under-stand their partitioning during volatilisation/condensation processes.There ismuch furtherwork to be done in this area.Moremeasurementsare needed of metals in both rock and aerosol in conjunction with gasmeasurements at diverse volcanoes to allow us to build up a morecomplete picture of metal volatility from silicate melts. Melt inclusionstudies, where possible, are also of great value (e.g., Edmonds, 2015).Further experiments to determine elemental volatilities from melts ofvarying composition will also undoubtedly play a role as well as ther-modynamic calculations and modelling (review of previous work inPokrovski et al., 2013).

6. Summary

We have learnt much about the impacts of volcanism on our planetby studying present-day activity. Studies during the satellite era haveshownus how large explosive eruptions can have clear impacts on plan-etary albedo, global atmospheric circulation patterns and stratosphericchemistry. Numerous more subtle and more poorly understood effects

Table 7Summary of volatile, moderately volatile and non-volatile elements as defined by themeasurements at Etna andMasaya (Table 6) compared to literature values for Tolbackik, Erte Ale andKudryavy.

Etna1 Masaya1 Tolbackik2 Erte Ale3 Kudryavy3,4

Volatile (volatilitycoefficient N1)

Bi, Hg, Re, Sb, Se, Tl Bi, Cd, Te, Tl As, Bi, Cd, Re, Se, Te, Tl Bi, Cd, Re, Se, Te, Tl 5B, Bi, Re, Se, Te

Moderately volatile(volatilitycoefficient 0.1–1)

As, Cd, Cu, In, Pb, Pt Cs, Cu, Pb, Sb Au, In, Pb, Pt Ag, As, Au, In, Pb, Pt, Sn As, Au, Cd, Mo, Sb, Tl,W

Non-volatile(volatilitycoefficient b0.1)

Ag, Al, Au, Ba, Ca, Co, Cr, Cs, Fe, Ga, Ge, Hf, K, 5Li,Mg, Mn, Mo, Na, Nb, Ni, Rb, Sc, Si, Sn, Sr, Te, Th, Ti,U, W, Y, Zn, Zr

Ba, Be, Co, Hf,Rb, Sr, Th, U, Y,Zn, Zr

Ag, 5B, Ba, Cs, Co, Cu, Fe, K,5Li, Mn, Mo, Na, Rb, Sb, Si,Sn, W, Zn

5B, Ba, Co, Cs, Cu, Fe, K,5Li, Mn, Mo, Na, Rb, Sb,Si, Zn, W

Ba, Co, Cs, Cu, Fe, K, 5Li,Mn, Na, Pb, Pt, Rb, Si,Sn, Zn

1 See Table 6.2 Estimated in Zelenski et al. (2014) who were sampling high-concentration volcanic gases from lava tube windows ~300 m away from the lava lake.3 Read off Fig. 11 in Zelenski et al. (2014) from data from Zelenski et al. (2013) for the basaltic Erta Ale volcano (collected in January 2011 from high-temperature (1084 °C) fumarolic

gases emitted from a spatter cone in themiddle of the north pit crater of the Erta Ale volcano) and Taran et al. (1995) for the basaltic andesite Kudryavy (fumarole condensates 585–940 °Ccollected during quiescent outgassing).

4 The vapour/melt partitioning coefficients calculated from fumarole condensates and associated lavas/rocks from Kudryavy by Pokrovski et al. (2013) from data presented in Taranet al. (1995) give Cl, Re, Se, B, S, Au, Ag, and Pt in the volatile category (N1), As, Mo and Sn in the moderately volatile category (0.1–1) and Pb, Sb, Zn, Zr, K, Cu, Na, Si and Fe in the non-volatile category (b0.1).

5 The results for B and Li are broadly consistentwith the findings of Edmonds (2015) based on Kīlaueamelt inclusions (vapour-melt partition coefficient for Li 0.02–0.24 and for B 0.014–0.5)other than in the case of B at Kudryavy. These data Kīlauea suggest that partitioning is inhibited at higher temperatures and depends also on the othermelt volatile phases. The co-depletion of Cland H2O in the M.I.s and the Fo contents of the host olivines supports low-pressure, high-temperature (N1100 °C) degassing.

174 T.A. Mather / Journal of Volcanology and Geothermal Research 304 (2015) 160–179

on the troposphere, such as modification of cloud properties, levels ofatmospheric oxidants, local sulphuric acid haze andmarine or terrestrialecosystems, have also been suggested following these large eruptions aswell as smaller explosive eruptions and passive degassing (summarisedin Fig. 1 and Table 3).

There are many areas where further measurements of present-dayvolcanic emissions have great potential to further our understandingof volcanism's impacts on the environment. Here I highlight 3 examples.

Firstly, while there is some convergence in published estimates forthe global SO2 flux from volcanoes (likely due at least in part to thedominant contribution from the highest flux volcanoes), more mea-surements are needed to better constrain the fluxes of other key volca-nic gases such as CO2 and halogens. On the scale of individual arcs evenSO2 fluxes are not well constrained over time (with a minimum uncer-tainty of a factor of 2 even for well studied arcs) andmuchmorework isneeded to determine whether or how key gas ratios (e.g., C/S) vary be-tween and along individual arcs. Even for high-temperature plumemeasurements C/S ratios vary by over 2 orders of magnitude globallywith significant variations within individual arcs and at individual vol-canoes. These uncertainties are among current barriers to advancingour understanding of volatile recycling through subduction zones.Variations in this recycling have been postulated to cause global changeduring periods of the Earth's past.

Secondly, we still lack a systematic understanding of volcanic halogeninputs to the stratosphere during explosive eruptions and their impacts

Fig. 5. (a) A plot of mean volatility coefficient based on the Etna andMasaya data presented in Tofmean volatility coefficient based on the Etna andMasaya data presented in Table 6 against thdisc would have condensed) calculated by Lodders (2003).

on ozone. Measurement-based estimates suggest a wide range of frac-tions of HCl reach the stratosphere, from~100% for the high latitude erup-tion of Hekla tomaybe b1% during the 2006 eruption from Soufrière HillsVolcano in the tropics (Table 4). Estimates from and input to modellingstudies cover a similar range. Further measurements and modelling areneeded in order to understand the scavenging processes operating in vol-canic plumes and how they vary with different plume and atmosphericconditions. Current measurements allow us to make a ‘best guess’ time-averaged HCl flux to the stratosphere of 0.006–0.24 Tg/yr. Recent studiesof volcanic Br emissions suggest that they could significantly impactstratospheric ozone levels following large eruptions.However,moremea-surements of halogens, especially Br, are also needed to develop a full un-derstanding of the potential past and future effects of volcanism onstratospheric ozone without the influence of anthropogenic halogenspecies.

Thirdly, the partitioning of trace metals between silicate melts andthe gas phase at the relatively low pressures characteristic of volcanicdegassing is poorly constrained. More comparison is needed of volatil-ities measured at different volcanic systems. I build on recent work toestimate the gas/melt partition coefficients using literature data fromEtna and Masaya and compare these with limited literature valuesfrom Tolbackik, Erte Ale and Kudryavy. This allows elements to begrouped into volatile/moderately volatile (As, Bi, Cd, Hg, Re, Se, Tl andIn) and non-volatile (Al, Ba, Ca, Co, Cr, Fe, Ga, Ge, Hf, K, Li, Mg, Mn, Na,Nb, Ni, Rb, Sc, Si, Sn, Sr, Th, Ti, U, Y, Zn and Zr). For some elements (Ag,

able 6 against the emanation coefficients (defined in Table 5) fromRubin (1997). (b) A plote condensation temperatures (at which 50% of an element's atoms in a protoplanetary gas

175T.A. Mather / Journal of Volcanology and Geothermal Research 304 (2015) 160–179

Au, B, Cs, Cu, Mo, Pb, Pt, Sb, Te and W) there is some disagreement intheir volatility between volcanic systems,which could reflect real differ-ences in the conditions at the different volcanoes or uncertainties in thedata. Comparison of these results with other systems to classify metalvolatility highlight broad consistency but also some potentially impor-tant differences.Moremeasurements in volcanic plumes and rock suitescombined with experimental studies and thermodynamic calculationshave the potential to further our systematic understanding of tracemetal volatility. Better understanding in this area has implications interms of metal release to the environment during volcanic activity aswell as their partitioning during other processes such as those involvedwith the evolution of magma geochemistry, ore formation and plane-tary evolution.

Acknowledgements

I have benefited from many useful and stimulating discussions thathave contributed in different ways to the writing of this review includ-ing those with Sandro Aiuppa, Lindy Elkins-Tanton, Frances Jenner,Charlie Langmuir, Rob Martin, Lawrence Percival, Fred Prata, DavidPyle, Anja Schmidt, Steve Self, Seb Watt, Jon Wade and Bernie Wood. Iwould like to thank the NERC (including grant reference NE/M000427/1) and the Leverhulme Trust for funding. I would also liketo thank the staff of St Anne's nursery for all their help and supportover the last 7 years.Morgan Jones, an anonymous reviewer and the ed-itor Lionel Wilson are thanked for their insightful comments on an ear-lier version of the manuscript. This review is dedicated to the memoryof Roland von Glasow and his legacy in terms of advancing our under-standing of the atmospheric chemistry occurring in volcanic plumes.

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