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Initiation of Snowball Earth with Volcanic Sulfur Aerosol Emissions F

Initiation of Snowball Earth with Volcanic Sulfur Aerosol Emissions F

Initiation of Snowball with volcanic sulfur aerosol emissions F. A. Macdonald1 & R. Wordsworth2, 1

1 Department of Earth and Planetary Sciences, Harvard University, 20 Oxford Street, Cambridge, Massachusetts 02138, USA 2 Harvard Paulson School of Engineering and Applied Sciences, Harvard University, 29 Oxford Street, Cambridge, Massachusetts 02138, USA

After 1.5 billion without significant glaciation, during the Period (720- 635 Ma), extended to the at least twice and glacigenic sediments were deposited on every major continent1,2. Although it is generally agreed that these events occurred via an ice- runaway3, the drivers of cooling past the critical bifurcation point are unknown. Geochronological constraints indicate that eruptions associated with the Franklin Large Igneous Province (LIP) coincided with the onset of the first of the Cryogenian Snowball Earth events, the Sturtian glaciation4. The Franklin LIP was the largest volcanic event5 and was emplaced into a sulfur evaporite basin at the equator4,6. Geochemical data on sills indicate wall- entrainment of sulfur7, which would have outgassed as SO2 and H2S and formed radiatively active sulfate aerosols. To constrain the magnitude and duration of an eruption needed to initiate a Snowball Earth, we performed coupled volcanic plume, aerosol and modeling. We show that in a cool climate with a low tropopause, multi- sulfur-rich fire-fountains at the equator can inject enough sulfur into the stratosphere to drive a runaway ice-albedo event. Coincidence has also been established between LIPs and mass extinctions8,9, but these events are variably associated with evidence for warming or cooling, and kill mechanisms are still unclear. Our results demonstrate that high eruptions, such as the Siberian Traps, which is associated with the Permo- extinction9, cause less cooling, or even rapid warming if they erupt over high albedo surfaces. Thus, we suggest that depending on pre-existing geological and climate conditions, natural sulfur aerosol emissions on decadal timescales can drive extreme , providing a link between LIPs, environmental destruction, and mass extinctions. Snowball Earth events represent the largest episodes of planetary climate change recorded in the geological record2, both on Earth and throughout the Solar , and yet we do not understand how or why these events occurred. It has been proposed that reduced volcanic 10 CO2 outgassing or increased planetary weatherability due to the predominance of low-latitude continents mantled with flood basalts led to a cool Neoproterozoic climate4,11-14, and that Snowball Earth was initiated by additional short-term perturbations to the 11,15 16-18 inventory or planetary albedo . Models invoking a short-term drawdown of CO2 or were motivated by an apparent relationship between perturbations to the and the onset of glaciation11,15. However, recent geochronology has demonstrated that the pre- Sturtian carbon excursion occurred more than 10 Myrs before the initiation of the by 716.5 ± 0.2 Ma4,13, whereas the eruption of the Franklin Large Igneous Province (LIP) was synchronous (Fig. 1). Volcanic rocks associated with the Franklin LIP cover an area > 3 Mkm2 over northern and southern (Extended data Fig. 1)5, and erupted at equatorial latitudes4,6. On Victoria Island of the Canadian Arctic, the Franklin LIP is represented by basalt and gabbroic dikes and sills that intruded , organic-rich shale, and sulfur-evaporite of the Shaler Supergroup19. Due to melting and assimilation of sulfur from the evaporites, many of the sills and dikes have extremely high sulfur concentrations, ranging from 100 to 100,000 ppm7, and sulfur isotope compositions indicative of contamination (Extended data Fig. 2). Importantly, these sulfur-rich sills were dated at 716.3 ± 0.5 Ma, within error of the Sturtian glaciation onset between 717.4 ± 0.2 Ma and 716.5 ± 0.2 Ma4 (Fig. 1). To evaluate the hypothesis that the emplacement of the Franklin LIP into a sulfur-rich basin liberated sulfate aerosols and initiated Snowball Earth, we first modeled the height to which sulfur-bearing volcanic plumes could be injected into the during the eruption of a LIP, the microphysical evolution of aerosols in the stratosphere, and the radiative effects of a large sulfur-rich eruption at low-latitude. Unlike the short-lived injection of sulfate aerosols to the stratosphere from the 1991 eruption of the Pinatubo stratovolcano, fire fountains associated with LIPs can erupt for several years and drive turbulent, volatile-rich convective plumes into the atmosphere20. We simulated the maximum height obtained by convective plumes using a steady- state model incorporating turbulent entrainment (see Extended Data). Plume height is a strong function of volume eruption rate, with values of 104-105 m3/s leading to plume heights greater than the present-day tropical tropopause (~12 km) under Neoproterozoic atmospheric conditions (Fig. 2). 21 Volcanic sulfur is outgassed as a combination of SO2 and H2S . Once in the atmosphere, these gases react with O2, H2O and OH to form H2SO4, which condenses with H2O onto condensation nuclei to form radiatively active sulfate aerosols22. Weak plumes inject sulfur into the troposphere, where sulfate aerosols are consumed in days to weeks and have little long-term radiative effect23. Conversely, strong plumes penetrate the tropopause and reach the stratosphere, 20 where H2SO4 aerosols can have lifetimes of a year or more . The primary determinant of tropopause height is surface temperature (Fig. 2a); a warmer surface injects more into the atmosphere, which reduces lapse rate. Hence a given eruption is more likely to penetrate the stratosphere when the background climate is cold20. This means that a pre-Snowball climate 4,11-13 where CO2 had already been drawn down to low levels would be more vulnerable to cooling via volcanic stratospheric aerosol injection. It also implies that once a volcanic eruptions begin, a can occur: surface cooling by aerosols lowers the tropopause height, making it easier for subsequent eruptions to penetrate the tropopause and recharge the aerosol content of the stratosphere. To evaluate the radiative forcing following stratospheric SO2 injection, we modeled sulfur chemistry and aerosol microphysics, incorporating particle growth, coagulation, sedimentation and mixing in a box model of the stratosphere. Radiative forcing is calculated using a 1D radiative-convective correlated-k model. We find that the radiative forcing caused by episodic injections of SO2 into the equatorial stratosphere increases with the quantity of SO2 injected (Fig. 3a), but the increase is sub-linear due to particle coagulation, which increases mean particle size and sedimentation rate for large eruptions24. For the case of episodic year-long 2 eruptions, mean global forcing increases from -2.7 W/m given 20 Gt SO2 injection (c.f. 25 2 Pinatubo ) to -12 W/m for 500 Mt SO2 (sensitivity to parameters described in Extended Data). Estimating the rate at which sulfur was released from the Franklin LIP on a decadal timescale is 26 challenging, but 500 Mt/yr is similar to estimates of SO2 release rate from more recent LIPs . The ability of volcanic aerosols to force a Snowball transition depends on the starting climate state and on the rate at which the mixed layer of the tropical ocean can cool. 3 Atmospheric CO2 levels in the Neoproterozoic are not well constrained but were probably low due to the predominance of equatorial continents1, which enhances weathering27. The amount of metamorphic CO2 released from the emplacement of gabbroic sills in the Franklin LIP was small relative to background levels and likely did not have a significant warming effect19. Moreover, relative to CO2, only small quantities of reflective aerosols are needed to alter Earth’s global radiative balance25. Based on radiative-convective modeling, we estimate a -10 W/m2 global mean aerosol radiative forcing would be sufficient to cause runaway glaciation for CO2 levels of 3000 ppm of below (Fig. 3b). For a representative mixed layer depth of 50 m, the cooling timescale is of order 3 years (see Extended Data). This timescale is compatible with geological observations: individual eruptions in recent mafic LIPs have been estimated to have lasted ~13 years with plume heights >15 km26. Larger flows, such as those associated with the Deccan and Siberian Traps, may have lasted an order of magnitude longer26 and had multiple eruptive centers. Hence a sequence of a few large, sulfur-rich eruptions of the type shown in Fig. 3A would have been sufficient to cause a Snowball transition (Fig. 3b). If the Franklin LIP caused a Snowball Earth, then why are other LIPs not also associated with Snowball Earth events? The Siberian Traps are similar in size to the Franklin LIP, but instead of cooling, they appear to be associated with warming and extinction28. Another LIP, the Central Magmatic Province (CAMP), may have caused a temporary glaciation followed by warming8. The different outcomes are likely related to several factors: the background climate conditions, changes in planetary albedo with different paleogeography, the latitude of the eruptions, the composition of the country rock that the LIPs were emplaced into (a sulfate-rich basin, coal deposits, and an ancient mountain belt for the Franklin, Siberian Traps and CAMP, respectively), and timing of SO2 injection relative to the cumulative release of greenhouse gases. Although Neoproterozoic paleogeography likely favored a cool climate1,27, the Siberian Traps and CAMP erupted during the ice-free Late and , respectively, when warm conditions extended to Earth’s polar regions29. Consequently, there was a higher ocean thermal inertia and higher mid-latitude tropopause that prevented stratospheric forcing by all but the most powerful volcanic eruptions (Fig. 2). Moreover, the Siberian Traps erupted at high latitude, whereas the CAMP and Franklin LIP erupted at equatorial latitudes6. Consequently albedo changes associated with the Siberian Traps would have primarily affected mid-to-high in one hemisphere, whereas the CAMP and Franklin LIP would have maximized albedo at low latitude, where solar forcing is highest. In general, because aerosol radiative forcing is highest given low surface albedo and a high latitude eruption in a cold climate occurs over snow or sea ice, an equatorial eruption is far more effective at forcing a Snowball transition than a high latitude one (see Extended Data). The geological record preserves evidence for two Cryogenian glaciations, the Sturtian and the Marinoan, and although there is not evidence for a large LIP during the onset of the , this may be a matter of preservation. The low-latitude rifting of continued throughout the Neoproterozoic, and it is likely that these were associated with additional LIPs. Geochronological constraints suggest the Sturtian glaciation lasted ~58 Myrs and that there was a short recurrence interval between the glaciations, between 8.6 and 19.4 Myrs13. If planetary weatherability remained high during the Cryogenian due to continued low- latitude paleogeography and recently emplaced continental flood basalt provinces14, transport- limited in the aftermath of the Sturtian glaciation may have returned the Earth on a 10 Myr timescale to a climate state sensitive to short-term perturbations in the concentration of greenhouse gases and albedo30. After the Marinoan glaciation, the removal of a basaltic carapace and drift of continents to higher latitudes1 likely reduced global weatherability and climate sensitivity. We argue that the proximal trigger for the Sturtian Snowball Earth was the emission of volcanic aerosols and a sudden increase in planetary albedo. Over the , the largest disturbances to surface environments coincided with LIPs and giant impacts. It appears that it was not the size of the LIP or the impactor that mattered most, but when and where it hit28. LIPs and meteorite impacts both hold the potential to rapidly change the environment through the release of metals, volatile gases and reflective particles, but critically it is the rate of change that determines whether species adapt or are exterminated. Although there has been much focus on the cumulative effects of volcanic emissions, we suggest that multi-year to decadal perturbations to albedo and climate from LIPs may instead drive the highest rate of environmental change. Thus, a better understanding of these formative events may require annual resolution of events that occurred hundreds of millions of years ago, which presents a substantial challenge for the geological community, but one that can potentially be met by the richness of the stratigraphic record.

1 Li, Z.-X., D.A.D., E. & Halverson, G. P. Neoproterozoic glaciations in a revised global palaeogeography from the breakup of Rodinia to the assembly of Gondwanaland. Sedimentary 294, 219-232 (2013). 2 Kirschvink, J. L. in The Biosphere (eds J. W. Schopf & C. Klein) 51-52 (Cambridge University Press, 1992). 3 Pierrehumbert, R., Abbot, D., Voigt, A. & Koll, D. Climate of the Neoproterozoic. Annual Review of Earth and Planetary Sciences 39, 417 (2011). 4 Macdonald, F. A. et al. Calibrating the Cryogenian. Science 327, 1241-1243 (2010). 5 Ernst, R. E. et al. Long-lived connection between southern Siberia and northern Laurentia in the Proterozoic. Geoscience 9, 464-469, doi:10.1038/NGEO2700 (2016). 6 Denyszyn, S. W., Halls, H. C., Davis, D. W. & Evans, D. A. D. and U- Pb geochronology of Franklin dykes in High Arctic Canada and : a revised age and paleomagnetic pole for constraining block rotations in the Nares Strait region. Canadian Journal of Earth Sciences 46, 689-705 (2009). 7 Bedard, J. H. et al. Geochemical database of Franklin sills, Natkusiak Basalts and Shaler Supergroup rocks, Victoria Island, Northwest Territories, and correlatives from Nunavut and the mainland. Geological Survey of Canada, Open File 8009 1, doi:10.4095/297842 (2016). 8 Schoene, B., Guex, J., Bartolini, A., Schaltegger, U. & Blackburn, T. J. Correlating the end-Triassic mass extinction and flood basalt at the 100 ka level. Geology 38, 387-390 (2010). 9 Burgess, S. D. & Bowring, S. A. High-precision geochronology confirms voluminous magmatism before, during, and after Earth’s most severe extinction. Science advances 1, e1500470 (2015). 10 McKenzie, N. R. et al. Continental arc volcanism as the principal driver of icehouse- greenhouse variability. Science 352, 444-447 (2016). 11 Schrag, D. P., Berner, R. A., Hoffman, P. F. & Halverson, G. P. On the initiation of snowball Earth. Geochemistry, , Geosystems 3, doi:10.1029/2001GC002991. (2002). 12 Godderis, Y. et al. The Sturtian 'snowball' glaciation: fire and ice. Earth and Planetary Science Letters 6648, 1-12 (2003). 13 Rooney, A. D. et al. Re-Os geochronology and coupled Os-Sr isotope constraints on the Sturtian snowball Earth. Proceedings of the National Academy of Science 1073, 1-6 (2014). 14 Cox, G. M. et al. A Basalt Weathering Trigger for Snowball Earth. Earth and Planetary Science Letters 446, 89-99 (2016). 15 Tziperman, E., Halevy, I., Johnston, D. T., Knoll, A. H. & Schrag, D. P. Biologically induced initiation of Neoproterozoic snowball-Earth events. Proceedings of the National Academy of Sciences 108, 15091-15096 (2011). 16 Bendtsen, J. & Bjerrum, C. J. Vulnerability of climate of Earth to sudden changes in insolation. Geophysical Research Letters 29 (2002). 17 Feulner, G., Hallmann, C. & Kienert, H. Snowball cooling after algal rise. Nature Geoscience 8, 659-662 (2015). 18 Stern, R. J., Avigad, D., Miller, N. & Beyth, M. in Links Between Geological Processes, Microbial Activities&Evolution of 313-337 (Springer, 2008). 19 Nabelek, P. I., Bédard, J. H. & Rainbird, R. H. Numerical constraints on degassing of metamorphic CO2 during the Neoproterozoic Franklin large igneous event, Arctic Canada. Geological Society of America Bulletin 126, 759-772 (2014). 20 Glaze, L. S., Self, S., Schmidt, A. & Hunter, S. J. Assessing eruption column height in ancient flood basalt eruptions. Earth and Planetary Science Letters (2015). 21 Textor, C., Graf, H. F., Herzog, M. & Oberhuber, J. Injection of gases into the stratosphere by explosive volcanic eruptions. Journal of Geophysical Research: (1984–2012) 108 (2003). 22 Turco, R., Hamill, P., Toon, O., Whitten, R. & Kiang, C. A one-dimensional model describing aerosol formation and evolution in the stratosphere: I. Physical processes and mathematical analogs. Journal of the Atmospheric Sciences 36, 699-717 (1979). 23 Chin, M. et al. A global three‐dimensional model of tropospheric sulfate. Journal of Geophysical Research: Atmospheres (1984–2012) 101, 18667-18690 (1996). 24 Pinto, J. P., Turco, R. P. & Toon, O. B. Self‐limiting physical and chemical effects in volcanic eruption clouds. Journal of Geophysical Research: Atmospheres 94, 11165- 11174 (1989). 25 McCormick, M. P., Thomason, L. W. & Trepte, C. R. Atmospheric effects of the Mt Pinatubo eruption. Nature 373, 399-404 (1995). 26 Self, S., Schmidt, A. & Mather, T. Emplacement characteristics, time scales, and volcanic gas release rates of continental flood basalt eruptions on Earth. Geological Society of America Special Papers 505, SPE505-516 (2014). 27 Marshall, H., Walker, J. C. G. & Kuhn, W. Long-term climate change and the geochmical cycle of carbon. Journal of Geophysical Research 93, 791-801 (1988). 28 Ganino, C. & Arndt, N. T. Climate changes caused by degassing of sediments during the emplacement of large igneous provinces. Geology 37, 323-326 (2009). 29 Taylor, E. L., Taylor, T. N. & Cúneo, N. R. The present is not the key to the past: a polar forest from the Permian of . Science, 1675-1675 (1992). 30 Mills, B., Watson, A. J., Goldblatt, C., Boyle, R. & Lenton, T. M. Timing of Neoproterozoic glaciations linked to transport-limited global weathering. Nature Geoscience 4, 861-864 (2011).

Million Years Ago (Ma) A. 850 800 750 700 650 600 Neoproterozoic Cryogenian Gaskiers Marinoan

-5 0 5 10 No glaciation for >1.5 billion years Sturtian Glaciation C (‰) 13 B.

U-Pb CA-ID-TIMS zircon data

Onset C.

* sulfur-rich sill 6A

725 720 715 710 (Ma) Maximum age constraint on onset of Sturtian glaciation Syn-glacial age constraint on Sturtian glaciation Dated Franklin large igneous province dikes and sills Figure 1. Geochronological constraints on the onset of the Sturtian glaciation compared to the most precise date on sulfur rich sills of the Franklin Large Igneous Province. A) Neoproterozoic timeline with carbon isotope chemostratigraphy, modified from Cox et al.14. B) Geochronological constraints on the onset of the Sturtian glaciation. C) Geochronological constraints on the Franklin large igneous province. Geochronological data and sources are in Extended Data.

Figure 2. Maximum plume height as a function of volume eruption rate under Neoproterozoic insolation and atmospheric conditions. Asterisks indicate tropopause height for each simulation. Black, red, green and blue lines correspond to simulations where the background CO2 levels were 30, 300, 3000 and 30000 ppm, respectively. Grey box brackets maximum plausible volume eruption rate during formation of the Franklin LIP (see Extended Data).

Figure 3. The effect of volcanic sulfate aerosol emission on global climate. A) Radiative effect of yearly equatorial volcanic eruptions under Neoproterozoic conditions. Black, red and green lines show simulations assuming 20, 100 and 500 Mt of SO2 injection to the stratosphere. Annual global mean forcings for each case are displayed. B) Bifurcation diagram of global mean temperature versus radiative forcing defined as the difference between outgoing longwave radiation and absorbed solar radiation. Dashed lines are unstable states. Starting from a mean surface temperature close to present-day (red star), a negative radiative forcing of < -10 W/m2 is sufficient to push the Neoproterozoic Earth (green line) into a Snowball state (blue star). Starting from a colder initial state, the required radiative forcing is lower. Supplementary Material: Initiation of Neoproterozoic Snowball Earth with volcanic sulfur aerosol emissions

F. Macdonald and R. Wordsworth

August 3, 2016

1 Materials and Methods

1.1 Geological Setting Igneous rocks associated with the Franklin LIP are dominated by mafic dikes and sills that intrude the Canadian Shield and Proterozoic basins on the Arctic margin of Laurentia (Fig. S1), and may be correlative with similar aged intrusions on the southern margin of Siberia (1). The dikes presumably fed large basaltic provinces that have since been largely eroded. Preserved extrusive volcanic rocks include the Natkusiak Formation basalt of Minto Inlier, Victoria Island (2, 3), the Mount Harper Volcanics of the Coal Creek inlier in the Ogilvie Mountains (4, 5), Yukon, the Pleasant Creek Volcanics of the Tatonduk inlier, which straddles the Yukon-Alaska border (6), and the Kikiktak Volcanics of Arctic Alaska (7). These extrusive rocks are composed predominantly of tholeitic basalt to basaltic andesites, typical of continental flood basalts. The Natkusiak magmatic assemblage consists of dolerite sills that intruded the Shaler Super- group (8, 9), northwesterly striking dikes, and a succession of plateau basalts (2). Sedimentary strata of the underlying Shaler Supergroup were updomed, beveled and faulted before and during the emplacement of the Natkusiak magmatic assemblage (9,10). Sills range in thickness from 5 to 100 m; some are differentiated with olivine accumulation at their base and pegmatites in their up- per levels (2). Exposures of the Natkusiak Formation basalts have been separated into a northern lobe (3) and a southern lobe (11) of exposures preserved in synclines. The Natkusiak Formation basalts are thickest in the northern lobe, with a maximum preserved thickness of ∼ 1100 m, and has been divided into seven members (3). In the southern lobe, the Natkusiak Formation is up to 200 m thick and has been separated into four units that have been correlated with Jefferson et al.’s (1985) ‘basal member’, ‘pyroclastic member’, and ‘lower massive member in the northern lobe (11). Thus, it appears that units above the ‘Lower Massive Member’ in the northern lobe are not preserved in the southern lobe.

Nabelek et al. (12) estimated the amount of metamorphic CO2 released from the emplacement of the gabbroic sills associated with the Franklin LIP into carbonate strata of the Shaler Supergroup. Along with carbonate, the Shaler Supergroup includes organic rich shale and evaporite deposits at multiple horizons that are several tens of meters thick. Evans et al. (13) estimated a volume of 30,000 km3 in the upper evaporites of the Kilian-Redstone River deposits and 90,000 km3 in the lower evaporites of the Minto Inlet Formation. Although the total amount and rate of S release during the emplacement of sills in the Shaler Supergroup is poorly constrained, disseminated copper sulfides are present within basalt flows, outlining clasts in the pyroclastic member, and as veins

1 in all of the members (3). Additionally, many of the sills and dikes have extremely high sulfur concentrations (14, 15). Although dikes have been observed that feed the sills, feeder dikes for the basalts have not been observed. Yet, it is assumed that at least some of the dikes fed the basalts as is suggested by their similar chemistry (2, 16). Some of the sills are folded around a volcanic center near Kilian Lake that has been correlated with the pyroclastic member (14), suggesting that at least some of the sills formed before or very early in the eruption sequence. However, other sills are vesiculated and feed flows higher in the basaltic stack. Although fissure eruption centers have not been identified associated with the Franklin, the exposure is limited and it does not mean they were not there, and several circular to keyhole shaped eruption centers have been mapped [the largest one identified is ∼ 3 × 2 km (14)]. Paleomagnetic data from the Franklin LIP provide the most robust Neoproterozoic pole for Laurentia and is an important tie point for Neoproterozoic paleogeography (17). Data spanning six key regions, the Canadian Mainland, Victoria Island, Baffin Island, Devon Island, Ellesmere Island and northern Greenland, yield a grand mean pole that restores Laurentia and the Natkusiak assemblage of the Franklin LIP at ca. 717 Ma to within 9 degrees of the equator (17,18) (Fig. S1). U-Pb zircon and baddeleyite ages from the Franklin LIP rocks span ca. 725-710 Ma (Table S1). This spread could represent distinct pulses of magmatism; however, for younger, better dated LIPs, most of the volume is erupted in << 1 Myr (19). Instead, this large spread may also be due to imprecise multigrain techniques and upper intercept ages, and we suggest that the most precise dates represent the best age constraint on the Franklin LIP. The most precise, single grain chemical abrasion-isotope dilution-thermal mass spectrometry (CA-ID-TIMS) on zircon and baddeleyite associated with the Franklin LIP yield ages between 717 and 716 Ma (Table S1). It has been also proposed that the ca. 725 Ma Dovyren Intrusive Complex in the Lake Baikal region of Russia represents a component of the Franklin LIP in SW Siberia (1). If true it would provide evidence that Franklin magmatism initiated well before the onset of the Sturtian glaciation and covered an area beyond that defined by the radiating dyke swarm in north-central to northeastern Laurentia. However, the Siberian age is imprecise and hence this correlation remains speculative. A quartz-rich rhyolite flow in member D of the Mount Harper Volcanics was dated at 717.4 ± 0.2 Ma and a tuff interfingering with glacial deposits of the overlying Rapitan Group was dated at 716.5 ± 0.2 Ma, both with U-Pb CA-ID TIMS on zircon (20). These dates have been interpreted to bracket the onset of the Sturtian glaciation (20). Additional recent ages on the Franklin LIP and the onset of the Sturtian glaciation confirm this age model (Table S1). In China, an ash bed within stratified of the lowermost Chang-an Formation was dated with secondary ion mass spectrometry (SIMS) on zircon at 715.9 ± 2.8 Ma (21). In Arctic Alaska, a 719.5 ± 0.3 Ma CA-ID-TIMS age on zircon from volcaniclastic strata at the top of the Kikiktak Volcanics and directly below the Hula Hula diamictite is consistent with both a Sturtian-age of the Hula Hula diamictite and a tight temporal relationship between the Franklin LIP and the onset of the Sturtian glaciation (7). Sills that intrude evaporite rich-strata of the upper Shaler Supergroup dated at 716.3 ± 0.5 Ma (ID-TIMS on baddelyite) (20) are indistinguishable from the constraints on the onset of the Sturtian glaciation (Fig. 1 in the main text). Importantly, the dated sill, sill 6a near Kilim Lake, is one of most sulfur-rich sills analysed to date (3).

1.2 Geochemistry Subaerial continental flood basalt (CFB) volcanism is characterized by the repeated effusion of ∼ 103 km3 of basaltic magma in less than 1 Myr. Each eruption comes from a dyke-fed fissure and

2 can last up to a decade or more (19). During eruptions, gas fluxes of > 1 Gt per year of SO2 and CO2 are possible (22). These volumes can increase when magmatic systems invade volatile-rich sediments as more volatiles are entrained into the melt. The occurrence of spatter, spatter-fed lava, and scoria mounds along eruptive fissures in the Columbia River CFB and in Icelandic fissure eruptions suggests the occurrence of violent fire-fountaining events (23). It has been estimated that these fire fountains exceeded 1.5 km in height with convective plumes rising to heights in excess of 13 km (22). Estimates of gas-release budgets with melt inclusion studies are limited to individual CFB eruptions in the mid- Columbia River CFB (24), the Deccan Traps (22), Siberian Traps (25), and studies of Icelandic fissure eruptions (26). These studies suggest that ∼75% of volative S species present in the rising melt are released at the vents largely as SO2 (22). The months-long Laki eruption in is estimated to have released ∼120 Mt of SO2 with maximum fluxes reaching 6 Mt per day (26). Estimation of S gas release from ancient eruptions is compromised by weathering and ground water alteration, which leads to the removal of volatile elements. Nonetheless, using melt inclusions it has been estimated that a single flow with in the Columbia CFB released ∼12

Gt of SO2 (24). Melt inclusion studies are beyond the scope of this study and it is not clear that there is material preserved in the Franklin LIP that is pristine enough for such a study. In crystal-poor or altered CFBs, original S content of the lava can be estimated with the S content of undegassed dikes and sills normalized to TiO2/FeO2 (27). S saturated melts from LIPs typically have between 1000 and 2500 ppm of S prior to degassing and flows contain much lower concentrations between 200 and 300 ppm (22). Major, minor, and trace element concentrations including S concentrations and S from sills and dikes associated with the Franklin LIP were compiled by Bedard et al. (15). S concentrations in sills range from 100-100,000 ppm and S isotopes show increasing variability with increasing concentration indicative of wall-rock assimila- tion (Fig. S2). At these high S concentrations, and with the late addition of S out of equilibrium with the melt, normalization of S concentrations to TiO2/FeO2 is both insignificant and not ap- propriate. If we assume ∼75% of S remains in melt (22), and that the sills were variably degassed from original S concentrations less than 2500 ppm, the upper end of this range is greater than that of basaltic flows in the Columbia River CFB (24), the Deccan Traps (22), Siberian Traps (25), and the Icelandic fissure eruptions (26). Crustal S is incorporated into mafic magmas predominantly via direct melting and assimilation of wall rock and xenoliths rather than by the liberation of sulfur during contact metamorphism in the thermal aureole (28). Although previous work has questioned the viability of releasing sulfur via intrusions into evaporites due to the high melting temperature of pure anhydrite (∼1450◦C), the melting temperature of mixtures of anhydrite and , which is probably closer to the com- position of the wall rock intruded by the Natkusiak sills, is much lower (∼1000◦C) (28). Moreover, hydrothermal systems associated with the Franklin LIP may have partially dissolved the evapor- ites, and if these fluids were saline, the solubility of anhydrite greatly increases with NaCl activity at temperatures between 600◦ and 800◦C(29). We suggest that excess S was incorporated into magmas primarily through the melting and dissolution of evaporite layers in the Minto Inlier and Killian formations of the Shaler Supergroup when intruded by sills, and then was rapidly degassed as bubbles in shallow intrusions, fissure eruptions, and in flows that were fed by these sills. Petrographic analysis has revealed abundant disseminated pyrite and immiscible metallic sulfur in the sills. Bulk S isotope studies preserve a wide range of compositions (Fig. S2), differing from the value of mantle derived LIPs (∼ 0‰ VCDT). Although Jefferson et al. (14) argued that trace

3 element compositions suggest that the sulfur was sourced from sulfidic shale of the Wynniatt Fm, the high S concentrations and the enriched S isotopic values of several of the sills are consistent with the S also coming from the entrainment of carbonate and sulfur evaporates into the sills (Fig. S2).

However, because we are concerned with decadal fluxes of SO2, the source and the total amount of S entrained and released in magmatic system are not critical. What is important is that the sills that are coincident with the onset of the Sturtian glaciation contain considerably more S than typical CFBs.

1.3 Radiative-convective climate model To estimate the climate effect of Franklin LIP eruptions, we combine radiative-convective, aerosol microphysics and volcanic plume modeling tools. In the following sections, each of these components is described in detail. To calculate Neoproterozoic tropopause height, we used a 1D radiative-convective model. The model atmospheric temperature structure is calculated as a function of surface temperature as- suming adjustment to a moist adiabat in the troposphere, following the methodology in (30). The tropopause height is then defined using the standard Earth-based empirical measure as where the magnitude of the vertical temperature gradient drops below 2 K/km (dT/dz > −2 K/km).

To relate CO2 molar concentration to surface temperature, we performed iterative 1D radiative- convective calculations with moist adjustment in the troposphere and an assumed relative humidity of 0.77 (31). In the stratosphere, the temperature profile is allowed to evolve freely until thermal equilibrium is reached. Atmospheric radiative transfer is calculated using a correlated-k approach (32). The absorption effects of CO2 and H2O (infrared and visible) and O3 (UV) are taken into account; absorption by other gases is set to zero. Rayleigh scattering in the shortwave by N2 and O2 is taken into account, but the radiative effects of clouds are neglected. The surface albedo is set to 0.24, a value that allows the model to reproduce the present-day mean surface temperature (288 K) in simulations with present-day solar flux, surface pressure and atmospheric composition. Such an elevated surface albedo implicitly accounts for the effect of low-lying clouds (a standard assumption in 1D radiative-convective modelling). The radiative forcing due to clouds is uncertain in Earth’s paleoclimate (33,34), so inclusion of a fixed cloud forcing here would not have led to greater insight.

For the CO2 and H2O line absorption, we computed high resolution spectra over a range of tem- peratures, pressures and gas mixing ratios. We used an 6 × 8 × 8 temperature, pressure and H2O molar concentration grid with values T = {100, 150,..., 350} K, p = {10−3, 10−2,..., 104} mbar and f = {10−8, 10−7,..., 10−1}, respectively. The correlated-k method was used to produce fast H2O coefficients for the radiative-convective simulations, with the CO2 mixing ratio chosen in advance for each calculation. The H2O continuum was implemented using the MT-CKD parametriza- tion (35, 36). We used 80 spectral bands in the longwave and 82 in the shortwave, and sixteen points for the g-space integration, where g is the cumulated distribution function of the absorp- tion data for each band. Twelve new bands were added to the shortwave compared to previous studies (37) to give sufficient resolution in the ultraviolet to model absorption by ozone accurately.

Ultraviolet O3 absorption cross-sections are derived from the MPI-Mainz UV/VIS Spectral Atlas (38) while the solar spectrum is taken from the VPL database (39). Calculations were performed with two O3 molar concentration profiles: one assumed present-day values, while the th other assumed an O3 profile appropriate to an atmosphere with levels 1/10 of today.

4 1.4 Aerosol radiative transfer We use Mie theory to calculate aerosol extinction efficiency, single scattering albedo and asym- metry parameter as a function of wavelength for a given particle size distribution ouput by the microphysical model. The aerosol radiative forcing is calculated using the two-stream approxima- tion, with Eddington scaling used in the shortwave and hemi-isotropic emission assumed in the longwave (40). We assume that absorption dominates in the longwave and that scattering dom- inates in the shortwave. Delta scaling is applied to the shortwave optical coefficients to ensure that the forward peak of the phase function is correctly captured at large particle size parameters. We describe the phase function using the Henyey-Greenstein approximation (41). For the aerosol particle refractive indices, we use the data given in (42), which were measured at stratospheric temperatures (∼200 K). Solar zenith angle θz was calculated by diurnal and seasonal averaging across the microphysics box domain, with weighting by the instantaneous insolation (43). For the ◦ equatorial eruption simulations (see next section), this yielded θz = 43.8 , while for the midlatitude ◦ case the value was θz = 55.4 . Finally, for the equatorial eruption simulations shown in Fig. 3A in the main text, we used a pre-aerosol tropical planetary albedo of 0.24. This was obtained by averaging albedo data from the Earth Radiation Budget Experiment (ERBE) mission (36,44) over the simulation domain. The same procedure was used to obtain the albedo value of 0.365 for the second mid-latitude eruption plot in Fig. S11.

1.5 Volcanic plume model The volcanic plume model simultaneously solves the equations of mass, momentum and energy transport for an axisymmetric, steady-state circular plume that entrains the surrounding air as it rises. These equations are defined as df d 1 = (ρ ur2) = 2αρ ur (1) dz dz g a df d 2 = (f u) = g(ρ − ρ )r2 (2) dz dz 1 a g df d 3 = (f c T ) = 2αρ urc T − ρ ur2g. (3) dz dz 1 p,g g a p,a a a with

2 f1 = ρgur (4)

f2 = f1u (5)

f3 = f1cp,gTg. (6)

Here α is an entrainment coefficient, z is altitude, u is the vertical speed of the plume, r is the plume radius, ρa is the density of ambient air, ρg is the mean density of gas in the plume, g is gravity, Ta is the local atmospheric temperature, Tg is the plume gas temperature and cp,a and cp,g are the specific heat capacity at constant pressure of ambient air and of the plume gas, respectively. Following (45), the plume specific heat capacity and gas constant are modified as it moves upwards to account for the entrainment of ambient air. After boundary conditions at the surface are chosen (see below), (1-3) are integrated upwards until the vertical speed of the plume equals zero. The height at which this occurs is then defined as the maximum plume height. This plume model is based on the methodology of (45). It is a

5 generalization of the classic incompressible model of Morton et al. (46), which has been validated via numerous laboratory experiments and comparison with observations of real volcanic eruptions (47). For the numerical integration upwards we use a simple forward Euler approach, with vertical resolution increased to the point where the results converge (we use 1000 levels based on trial and error). To initialize our model, we use the same approach as in a previous study that successfully predicted the plume heights of the historic Laki, Mauna Loa and Askja eruptions (47). Specifically, 3 we link Q to the volume eruption rate Verupt [m /s] via the expression

Q = ρmagmaVerupt[(1 − qf − qv)cp,magma∆Tc + (qf cp,magma + qvcp,v)∆Tg]. (7)

Equation (8) assumes that the plume is heated by a combination of volatiles, fines and clasts, with a negligible contribution from fire fountain lava blobs. Here ρmagma is the magma density, 3 which we take to be 2900 kg/m — a value appropriate to basaltic composition, qf is the specific concentration of fines [kg/kg] and qv is the specific concentration of volatiles. cp,magma and cp,v are the specific heat of magma and volatiles (the latter is assumed to be dominated by water vapor), respectively. ∆Tc is the temperature drop of clasts in the fire fountain while ∆Tg is the temperature drop of volatiles and fines in the plume. See Table S2 for definitions and values of all model parameters. We assume conservative values for the key unknown parameters: qf = 0.01 kg/kg, qv = 0.02 kg/kg, ∆Tc = 50 K and ∆Tg = Tg,0 − Tg,1 = 1400 − 220 = 1180 K (here Tg,1 is the as- sumed final plume temperature). We assume the dominant volatile by mass is water vapor. The contribution of sulfur and other comparatively minor species to the heat budget is neglected. The initial mass flux is calculated from the heat flux assuming a circular vent as (45)

Q f1,0 = . (8) πcp,v(Tg,0 − Ta,0) where 0 denotes initial (surface) variables. Fissure systems tend to produce lines of circular cones to key-hole shaped vents that range from tens of meters to kilometers in diameter, but more elongated eruptive centers also occur. Glaze et al. (2015) (48) found that elongated geometries produce larger plumes, but for simplicity we use a conservative circular geometry here.

1.6 Aerosol microphysical model We model key aspects of the sulfate aerosol microphysics, focusing on those processes that play the most important role in determining total radiative forcing. Our model describes the evolution of

SO2 and H2SO4 gas, Aitken particles (aerosol condensation nuclei), and the aerosols themselves. Aerosol particles undergo nucleation onto Aitken particles, growth, sedimentation and coagulation. In the microphysical model, the stratosphere is treated as a single box that exchanges material with the troposphere. The advantage of this idealized approach is that it captures the key processes governing aerosol optical depth and mean radius while allowing rapid exploration of sensitivity to poorly constrained parameters. Given the uncertainties, this approach is a more robust one than using a complex off-the-shelf model tuned to present-day Earth conditions. Once they reach the tropical stratosphere, sub-tropical transport barriers cause aerosols lofted there to remain mostly confined to that region (49, 50). Orbital observations of stratospheric aerosol loading after Pinatubo showed that aerosol optical depth at 0.5 µm remained above 1/2

6 the peak value for around 45 weeks, with a strong peak in forcing in the 30◦S-30◦N latitude region (51, 52). Modeling the stratospheric meridional circulation in the Neoproterozoic is challenging due to uncertainties in topographic wave forcing and stratospheric wave-mean interaction, but the aerosol transport following eruptions of the Franklin LIP should have been broadly similar. In this study we treat the tropical and midlatitude stratosphere as distinct domains. For the tropical eruption simulations shown in the main text, we defined the domain base as 16 km and the domain volume as V = 2πrEs∆z, with rE Earth’s radius, s = 1.4rE, corresponding to a latitudinal extent from 30◦ S to 30◦ N, and ∆z = 15 km the height of the simulation domain. Our microphysical model is sectional, with 82 particle bins ranging from 0.01 µm to 4.14 µm radius. Bin spacing is defined geometrically, with the particle volume increasing by a factor of 1.25 each time the bin number increases. Once the microphysical model produces aerosol mass loading and size distribution data as a function of time, its results are fed into the radiative transfer model to yield net radiative forcing in the domain. Global net radiative forcing is then calculated by multiplying 2 the results by a factor 2πrEs/4πrE. The model prognostic equations take the general form

∂n˜k(t) ∂n˜k n˜k = Pk − Lk − − (9) ∂t ∂t growth τexchange

∂ni(ri, t) ∂ni ∂ni ni ni = + − − (10) ∂t ∂t growth ∂t coag τsedim τexchange wheren ˜k and ni are molar concentrations of gas and aerosol species, respectively, and ri is the aerosol particle size in bin i. Pk and Lk are chemical production and loss terms, while the time derivative terms describe changes due to particle growth and coagulation. τexchange and τsedim are timescales for exchange of material with the troposphere and adjacent stratospheric regions, and for aerosol sedimentation. The equations are solved using the Klopfenstein-Shampine algorithm (ode15s in Matlab).

The transformation of SO2 into H2SO4 in the stratosphere proceeds by a chain of photochemical reactions, the most important of which are (53)

SO2 + OH + M → HSO3 + M (11)

HSO3 + OH → SO3 + H2O (12)

SO3 + H2O → H2SO4 (13)

In our model, the chemical production and loss terms Pi and Li are defined by the key sulfur reactions shown in Table S3. We neglect the outgassing of H2S in our model. As we find that the SO2 injection rates to the stratosphere required to force a Snowball transition are orders of magnitude lower than the upper limits on Franklin SO2 outgassing rates, inclusion of H2S would have added complexity to the model without fundamentally changing the nature of the conclusions. The model is initialized by the injection of a given mass of SO gas m into the box. Injected 2 SO2 masses are calculated as multiples of the SO injected by Pinatubo [approx. m = 20 Mt; (19)]. 2 SO2,0 The number density of H2O and the background gas density [M] are calculated from a standard present-day atmospheric profile (54) (Table S2), while the OH number density is set to yield the characteristic SO2 to H2SO4 conversion timescale observed after Pinatubo of 35 days (55). The sensitivity of the results to changes in OH levels are described in the next section. At the first

7 model timestep, the aerosol distribution is set to consist of condensation nuclei only. We assume that condensation nuclei are transported into the stratosphere along with the sulfur-bearing gas. The number concentration in the first particle bin is initialized as n0 = n (m /m ), where 1 Cn SO2 SO2,0 m = 20 Mt. We find that n = 100 particles cm−3 yields closest correspondence with SO2,0 Cn Pinatubo observations. This value is similar to the observed condensation nuclei number density at the tropopause on present-day Earth (56). Sensitivity of the results to variations in this parameter are also described in the next section.

Aerosol particle nucleation and growth

Aerosol particles grow through absorption of both H2SO4 and H2O. Equilibration with H2O is rapid, however, and the variation of equilibrium H2O:H2SO4 droplet weight ratios with height in the atmosphere is not high (53). For simplicity, here we assume that droplets have a fixed H SO specific concentration q = 0.75 kg/kg, a value appropriate to Earth’s present-day 2 4 H2SO4 lower stratosphere (53). Aerosol particles are assumed to form on Aitken particles, which we treat as a monodisperse population of small (rp = 0.01 µm) particles. The growth of particles by condensation of H2SO4 is given by dr mH SO DH SO (˜nH SO − n˜sat,H SO ) g = p = 2 4 2 4 2 4 2 4 (14) dt r ρ q (1 + λ Kn ) p p H2SO4 p 1 where r is the particle radius, ρ is the particle density, m is the molecular mass of H SO , p d H2SO4 2 4 D is the molecular diffusion coefficient of H SO in air (calculated following the procedure for H2SO4 2 4 trace gases described in (57)), n is the number density of H SO molecules near the droplet H2SO4 2 4 and n is the H SO condensation number density (53). The latter is derived from the sat,H2SO4 2 4 H2SO4 vapor pressure curve, for which we use the coefficients given in (58). Growth rate was then converted to the rate of change of aerosol number concentration by a derivative with respect to particle radius

∂ni ∂ = (gni) . (15) ∂t growth ∂r Kn = l /r is the Knusden number for H SO diffusion in air, while we define λ here as 1 H2SO4 p 2 4 p

1.33 + 0.71/Kn1 λp = . (16) 1 + 1/Kn1 The H SO diffusion coefficient D and mean free path in air l were calculated using 2 4 H2SO4 H2SO4 standard methods (57), with Sutherland’s formula    1.5 −5 416.16 Tstrat ηa = 1.8325 × 10 (17) Tstrat + 120 296.15 used to calculate the dynamic viscosity of air as a function of temperature.

Aerosol particle sedimentation Sedimentation is modeled using the Stokes velocity equation 2 v = f r2gρ /η , (18) sedim 9 C p p a

8 where g is gravity and ρa is air density. Here

−C/Kn2 fC = 1 + Kn2(A + Be ) (19) is the Stokes-Cunningham low pressure correction factor, with A = 1.249, B = 0.42, C = 0.87 , Kn2 = lair/rp the Knudsen number for air diffusion and lair the mean free path of air (57). The sedimentation timescale is then calculated as τsedim = zfall/vsedim, where zfall = 5.9 km is the distance that an average particle needs to fall to exit the stratosphere.

Aerosol particle coagulation When they collide, aerosol particles may coagulate. Over time this process shifts the particle distribution towards larger radii, increasing the sedimentation rate and decreasing the radiative forcing per unit mass. We describe coagulation using the discrete equation

i−1 ∂ni 1 X X = + Kj,i−jnjni−j − ni Ki,jnj (20) ∂t 2 coag j=1 j=1 where ni is the aerosol number concentration in the ith bin, t is time and Ki,j is the coagulation kernel. Total particle volume conservation is ensured by use of a bin-dependent volume fraction correction term (59). We validated our coagulation scheme by comparison with analytical results for the constant Ki,j = K0 case [Smoluchowski’s result; see e.g. (60)]. The coagulation kernel is calculated using the Fuchs equation, which is valid both in the high and low particle Knudsen number regimes. Specifically, for two particles of diameter dp,1 and dp,2, we define K12 = 2π(D1 + D2)(dp,1 + dp,2)β (21) with " #−1 d + d 8(D + D ) β = p,1 p,2 + 1 2 , (22) p 2 2 p 2 2 dp,1 + dp,2 + 2 g1 + g2 (c1 + c2)(dp,1 + dp,2)

1 h 3 2 2 3/2i gi = (dp,i + li) − (dp,i + li ) − dp,i, (23) 3dp,ili mean particle speed r 8kBT ci = , (24) πmi mean free path 8Di li = , (25) πci and particle diffusivity kBT fC Di = (26) 3πηadp,i with mi the particle mass and kB Boltzmann’s constant (60).

9 1.7 Bifurcation Plot Figure 3B in the main text was produced using a classical Budkyo-Sellers type model of the ice- albedo feedback. We assume that surface albedo depends on temperature as  Ai T ≤ Ti  2 (T −T0) A(T ) = A0 + (Ai − A0) (T −T )2 Ti < T < T0 (27)  i 0 A0 T ≥ T0 with T0 the temperature at which ice entirely disappears from the surface, Ti the temperature at which ice coverage is global, Ai the albedo of snow/ice and A0 the albedo of the ice-free planet. Justification of the terms in this equation are given in standard climate textbooks (36). Here we take representative values Ti = 260 K, T0 = 290 K and Ai = 0.6. A0 is set to be 0.24, as in the ice-free tropopause height simulations. The radiative-convective model was then run assuming a

1 bar N2 atmosphere and various CO2 molar concentrations to produce absorbed solar radiation and outgoing longwave radiation for a range of surface temperatures Ts. The radiative imbalance ∆F = OLR − ASR was then produced as a function of Ts via interpolation. In some coupled 3D general circulation models with ocean dynamics included, Neoproterozoic

Snowball glaciations occur close to lower CO2 partial pressures than those modeled here (61–63). While we do not attempt to capture the effects of sea ice and ocean dynamics here, we note that the basic effects of aerosol forcing would play out similarly starting from a lower CO2 climate state. As discussed in the main text, CO2 levels, while not well constrained prior to glacial onset, were probably low due to the predominance of low-latitude continents (64) and enhanced global weatherability (65).

1.8 Ocean Thermal Inertia The energy balance equation for the tropical ocean mixed layer can be approximated as d (ρ c T d) = ASR − OLR = ∆F (28) dt w p,w

3 with ρw = 1000 kg/m the density of , cp,w = 4200 J/kg/K the heat capacity of seawater, T temperature and d the depth of the mixed layer (36). For a radiation imbalance ∆F caused by aerosol forcing we can write the characteristic cooling timescale as ρ c d∆T τ = w p,w (29) cool ∆F Given ∆F = 10 W/m2, a required global mean ∆T of 5 K and a mixed layer depth of 50 m,

ρ c d∆T τ = w p,w = 3.3 y. (30) cool ∆F While certainly missing many additional effects that could be resolved in a 3D coupled general circulation model, this simple calculation demonstrates that the radiative forcing produced by SO2 emissions from volcanism over 3 years or more is capable of tipping Earth into a Snowball state.

10 2 Model Validation and Sensitivity Tests

2.1 Radiative-convective climate model To validate our radiative-convective model, we ran it assuming present-day globally averaged inso- lation and surface pressure, clear-sky conditions and a solar zenith angle of 60◦. Results are shown in Figure S3. As can be seen, the model reproduces Earth’s present-day mean temperature profile reasonably accurately. Figure S4 shows the temperature profiles with varying atmospheric CO2 used to create Figure 2 in the main text.

2.2 Volcanic plume model We compared our plume model with analytic and numerical results from the literature (45–47). Figure S5 shows the results of a single simulation with moderate values of heat flux and initial plume velocity. This figure reproduces Fig. 2 in (45), validating our model. Figures S6-S7 show the results of the combined plume / tropopause height modelling for different atmospheric ozone and nitrogen abundances. These figures demonstrate that our conclusions are not dependent on uncertainties in these parameters.

2.3 Aerosol Microphysics In addition to testing subcomponents of the model vs. analytic results (see previous section), we validated the aerosol microphysical model by comparison with observations of the Pinatubo eruption event. Given a single injection of 20 Mt SO2 into the tropical stratosphere, our model predicts a global mean radiative forcing of -2.0 W/m2 over two years, which falls in the -1 to -3 W/m2 range observed in the two years after Pinatubo (51). Our predicted peak optical depth at 0.5 µm wavelength is 0.2 and our peak effective particle radius is around 0.5 µm, both of which are also close to the observed post-Pinatubo values (66). Finally, our peak aerosol sulfur mass burden is 8 Mt of S, which is close to the values obtained by complex 3D models with coupled microphysics and chemistry (Figure S8) (67). To test the response of the microphysical model to uncertain parameters, we also performed a range of sensitivity studies. See Table S4 for details. Figure S9 shows the effective particle radius (as defined in (68)) and stratospheric aerosol S burden as a function of time for the three simulations displayed in Figure 3A in the main text. Figure S11 shows the radiative forcing from a similar set of simulations run in a domain centered on the mid-latitude northern hemisphere, assuming an ice-covered surface. Radiative forcing is significantly less negative in this case due to the increased surface albedo, which decreases the aerosol cooling effect. Aerosol transport rates from mid-latitudes to the tropics in the stratosphere are slow, so in contrast to the Franklin LIP, a high-latitude eruption would be unlikely to force a Snowball transition. Finally, we also tested the dependence of our radiative forcing predictions on the pre-aerosol planetary albedo. Results are shown in Figs. S10 and S11.

11 References

1. R. Ernst, et al., Nature Geoscience in press, XXX (2016).

2. J. Dostal, W. Baragar, C. Dupuy, Canadian Journal of Earth Sciences 23, 622 (1986).

3. C. Jefferson, W. Nelson, R. Kirkham, J. Reedman, R. Scoates, Current Research Part A: Geological Survey of Canada Paper pp. 203–214 (1985).

4. G. Cox, et al., Mount Harper Volcanic Complex, Ogilvie Mountains: A far-flung occurrence of the Franklin Igneous Event (Yukon Geological Survey, Whitehorse, Yukon, 2013), pp. 19–36.

5. C. F. Roots, Regional tectonic setting and evolution of the Late Proterozoic Mount Harper volcanic complex, Ogilvie Mountains, Yukon, Thesis, Carleton University (1987).

6. F. Macdonald, et al., Neoproterozoic and early correlations in the western Ogilvie Moutnains, Yukon (Yukon Geological Survey, Whitehorse, 2011), pp. 161–182.

7. G. Cox, et al., Lithosphere L435-1 (2015).

8. G. Young, The Amundsen Embayment, Northwest Territories: Relevance to the Upper Pro- terozoic evolution of North America (Geological Survey of Canada Paper 81-10, 1981), pp. 203–218.

9. R. H. Rainbird, Journal of Geology 101, 305 (1993).

10. J. Bedard, et al., Geological Society of America Bulletin 124, 723 (2012).

11. N. M. B. Williamson, L. Ootes, R. H. Rainbird, J. H. B´edard,B. Cousens, Bulletin of Vol- canology 78, 1 (2016).

12. P. I. Nabelek, J. H. B´edard,R. H. Rainbird, Geological Society of America Bulletin 126, 759 (2014).

13. D. A. Evans, Nature 444, 51 (2006).

14. C. Jefferson, et al., Geological Survey of Canada, Open File Report 2789 (1994).

15. J. Bedard, et al., Geological Survey of Canada, Open File 8009 1 (2016).

16. L. M. Heaman, A. N. LeCheminant, R. Rainbird, Earth and Planetary Science Letters 109, 117 (1992).

17. K. Buchan, et al., Tectonophysics 319, 167 (2000).

18. S. Denyszyn, D. Davis, H. Halls, Canadian Journal of Earth Sciences 46, 155 (2009).

19. S. Self, A. Schmidt, T. A. Mather, Geological Society of America Special Papers 505, SPE505 (2014).

20. F. Macdonald, et al., Science 327, 1241 (2010).

21. Z. Lan, et al., Research 255, 401 (2014).

12 22. S. Self, T. Thordarson, M. Widdowson, Elements 1, 283 (2005).

23. S. Self, M. Widdowson, T. Thordarson, A. E. Jay, Earth and Planetary Science Letters 248, 518 (2006).

24. T. Thordarson, S. Self, Journal of Volcanology and Geothermal Research 74, 49 (1996).

25. B. A. Black, L. T. Elkins-Tanton, M. C. Rowe, I. U. Peate, Earth and Planetary Science Letters 317, 363 (2012).

26. T. Thordarson, S. Self, N. Oskarsson, T. Hulsebosch, Bulletin of Volcanology 58, 205 (1996).

27. T. Thordarson, S. Self, D. Miller, G. Larsen, E. Vilmundard´ottir, Geological Society, London, Special Publications 213, 103 (2003).

28. J. Robertson, E. M. Ripley, S. J. Barnes, C. Li, Economic Geology 110, 1111 (2015).

29. R. C. Newton, C. E. Manning, Journal of Petrology 46, 701 (2005).

30. R. D. Wordsworth, R. T. Pierrehumbert, The Astrophysical Journal 778, 154 (2013).

31. S. Manabe, R. T. Wetherald, Journal of Atmospheric Sciences 24, 241 (1967).

32. R. Wordsworth, et al., Icarus 222, 1 (2013).

33. C. Goldblatt, K. Zahnle, Climate of the Past Discussions 6, 1163 (2010).

34. L. C. Sloan, D. K. Rea, Palaeogeography, Palaeoclimatology, Palaeoecology 119, 275 (1996).

35. S. A. Clough, F. X. Kneizys, R. W. Davies, Atmospheric Research 23, 229 (1989).

36. R. T. Pierrehumbert, Principles of Planetary Climate (Cambridge University Press, 2011).

37. R. Wordsworth, R. Pierrehumbert, Science 339, 64 (2013).

38. A. Serdyuchenko, V. Gorshelev, M. Weber, W. Chehade, J. P. Burrows, Atmospheric Measure- ment Techniques 7, 625 (2014).

39. A. Segura, et al., Astrobiology 3, 689 (2003).

40. W. E. Meador, W. R. Weaver, Journal of the Atmospheric Sciences 37, 630 (1980).

41. K. N. Liou, An introduction to atmospheric radiation, vol. 84 (Academic press, 2002).

42. R. T. Tisdale, D. L. Glandorf, M. A. Tolbert, O. B. Toon, Journal of Geophysical Research: Atmospheres 103, 25353 (1998).

43. D. L. Hartmann, Global physical climatology, vol. 103 (Newnes, 2015).

44. B. R. Barkstrom, Bulletin of the American Meteorological Society 65, 1170 (1984).

45. L. S. Glaze, S. M. Baloga, Journal of Geophysical Research: Atmospheres (1984–2012) 101, 1529 (1996).

13 46. B. R. Morton, G. Taylor, J. S. Turner, Proceedings of the Royal Society of London A: Mathe- matical, Physical and Engineering Sciences (The Royal Society, 1956), vol. 234, pp. 1–23.

47. R. B. Stothers, J. A. Wolff, S. Self, M. R. Rampino, Geophysical Research Letters 13, 725 (1986).

48. L. S. Glaze, S. Self, A. Schmidt, S. J. Hunter, Earth and Planetary Science Letters (2015).

49. C. M. Volk, et al., Science 272, 1763 (1996).

50. J. L. Neu, R. A. Plumb, Journal of Geophysical Research: Atmospheres 104, 19243 (1999).

51. M. P. McCormick, L. W. Thomason, C. R. Trepte, et al., Nature 373, 399 (1995).

52. W. B. Grant, et al., Journal of Geophysical Research: Atmospheres 101, 3973 (1996).

53. R. P. Turco, P. Hamill, O. B. Toon, R. C. Whitten, C. S. Kiang, Journal of the Atmospheric Sciences 36, 699 (1979).

54. G. P. Anderson, S. A. Clough, F. X. Kneizys, J. H. Chetwynd, E. P. Shettle, AFGL atmospheric constituent profiles (0-120 km), Tech. rep., DTIC Document (1986).

55. W. G. Read, L. Froidevaux, J. W. Waters, Geophysical research letters 20, 1299 (1993).

56. R. Turco, P. Hamill, O. Toon, R. Whitten, C. Kiang, NASA Technical Report (1979).

57. M. Z. Jacobson, Fundamentals of atmospheric modeling (Cambridge University Press, 2005).

58. M. Kulmala, A. Laaksonen, The Journal of Chemical Physics 93, 696 (1990).

59. M. Z. Jacobson, R. P. Turco, E. J. Jensen, O. B. Toon, Atmospheric Environment 28, 1327 (1994).

60. J. H. Seinfeld, S. N. Pandis, Atmospheric chemistry and physics: from air pollution to climate change (John Wiley & Sons, 2012).

61. A. Voigt, J. Marotzke, Climate dynamics 35, 887 (2010).

62. R. T. Pierrehumbert, D. S. Abbot, A. Voigt, D. Koll, Annual Review of Earth and Planetary Sciences 39, 417 (2011).

63. J. Yang, W. R. Peltier, Y. Hu, Climate of the Past 8, 907 (2012).

64. Z. X. Li, D. A. D. Evans, G. P. Halverson, Sedimentary Geology 294, 219 (2013).

65. H. Marshall, J. C. G. Walker, W. Kuhn, Journal of Geophysical Research 93, 791 (1988).

66. P. B. Russell, et al., Journal of Geophysical Research: Atmospheres 101, 18745 (1996).

67. S. S. Dhomse, et al., Atmospheric Chemistry and Physics 14, 11221 (2014).

68. J. E. Hansen, L. D. Travis, Space Sci. Rev. 16, 527 (1974).

69. G. J. Baldwin, E. C. Turner, B. S. Kamber, Precambrian Research 278, 303 (2016).

14 70. S. A. Bowring, J. P. Grotzinger, D. J. Condon, J. Ramezani, M. Newall, American Journal of Science 307, 1097 (2007).

71. Z. Lan, et al., Precambrian Research 255, 401 (2014).

72. Z. Lan, X.-H. Li, Q. Zhang, Q.-L. Li, Precambrian Research 267, 28 (2015).

73. F. Macdonald, et al., Science 327, 1241 (2010).

74. C. Fanning, P. K. Link, Geology 32, 881 (2004).

75. G. Cox, et al., Lithosphere L435-1 (2015).

76. R. P. Tollo, J. N. Aleinikoff, American Journal of Science 296, 1045 (1996).

77. X. S. Z. H.-B. C. Y. L. Z.-X. S. B. L. D.-Y. Z. C.-M. Y. X.-L. Xu, B., Precambrian Research 168, 247 (2009).

78. Q.-R. Zhang, X.-H. Li, L.-J. Feng, J. Huang, S. Biao, Journal of Geology 116, 423 (2008).

79. S. Denyszyn, D. Davis, H. Halls, Canadian Journal of Earth Sciences 46, 155 (2009).

80. S. Denyszyn, H. Halls, D. Davis, D. Evans, Canadian Journal of Earth Sciences 46, 689 (2009).

81. L. M. Heaman, A. N. LeCheminant, R. Rainbird, Earth and Planetary Science Letters 109, 117 (1992).

82. S. Pehrsson, K. Buchan, Canadian Journal of Earth Sciences 36, 65 (1999).

83. B. Schoene, J. Crowley, D. Condon, M. Schmitz, S. Bowring, Geochimica et Cosmochimica Acta 70, 426 (2006).

84. S. P. Sander, et al., Chemical kinetics and photochemical data for use in atmospheric studies evaluation number 15, Publication 06-2 , NASA JPL (2006).

85. T. Reiner, F. Arnold, The Journal of chemical physics 101, 7399 (1994).

15 Age constraints on Sturtian glaciation, techniques and data sources Location Age (Ma) (+) (-) Technique Mineral/Grains Formation Rel. Reference NWT, Laurentia 711.3 0.3 0.3 U-Pb CA-ID-TIMS detrital zircon Sayeuni syn Baldwin et al., 2016 Oman, Arabia 711.5 0.3 0.3 U-Pb CA-ID-TIMS magmatic zircon Ghubrah syn Bowring et al., 2007 South China 714.0 8.0 8.0 U-Pb SIMS magmatic zircon Liantuo max Lan et al., 2015 South China 715.9 2.8 2.8 U-Pb SIMS magmatic zircon Gongdong max Lan et al., 2014 South China 716.1 3.4 3.4 U-Pb SIMS magmatic zircon Gongdong max Lan et al., 2014 Yukon, Laurentia 716.5 0.2 0.2 U-Pb CA-ID-TIMS magmatic zircon Eagle Creek syn Macdonald et al., 2010 Idaho, Laurentia 717.0 4.0 4.0 U-Pb SHRIMP clast in diamictite Pocatello max Fanning & Link, 2004 Yukon, Laurentia 717.4 0.1 0.1 U-Pb CA-ID-TIMS magmatic zircon Mt. Harper Vol. max Macdonald et al., 2010 Arctic Alaska 719.5 0.3 0.3 U-Pb CA-ID-TIMS detrital zircon Kikitak Vol. max Cox et al., 2015 Virginia, Laurentia 722.0 3.0 3.0 U-Pb ID-TIMS bulk zircon Robinson River Vol. max Tollo & Aleinikoff, 1996 South China 724.0 12.0 12.0 U-Pb SHRIMP detrital zircon Liantuo max Gao & Zhang, 2009 Virginia, Laurentia 724.0 3.0 3.0 U-Pb ID-TIMS bulk zircon Robinson River Vol. max Tollo & Aleinikoff, 1996 Tarim, China 725.0 10.0 10.0 U-Pb SHRIMP magmatic zircon Bayisi max Xu et al., 2009 South China 725.0 10.0 10.0 U-Pb SHRIMP detrital zircon Banxi max Zhang et al., 2008

Age constraints on Franklin large igneous province, techniques and data sources Location Age (Ma) (+) (-) Technique Mineral/Grains Formation Rel. Reference

16 Devon Island 712 2 2 U-Pb ID-TIMS bulk baddeleyite Granville Fjord sill sill Denyszyn et al., 2009b Ellesmere Island 713 2 2 U-Pb ID-TIMS single grain baddeleyite Clarence head dike Denyszyn et al., 2009a Ellesmere Island 713 3 3 U-Pb ID-TIMS bulk baddeleyite Clarence head dike Denyszyn et al., 2009a Victoria Island 714.2* 2.6 2.5 upper-intercept bulk baddeleyite Low. Sills, Shaler Gp sill Heaman et al., 1992 Ellesmere Island 716 1 1 U-Pb ID-TIMS single grain baddeleyite Clarence head dike Denyszyn et al., 2009a Borden 716 4 5 U-Pb ID-TIMS bulk baddeleyite Borden dikes dike Pehrsson & Buchan, 1999 Victoria Island 716.3 0.5 0.5 U-Pb ID-TIMS single grain baddeleyite Sill 6a, Shaler Gp sill Macdonald et al., 2010 Bathhurst Inlet 720* 1 1 Pb-Pb ID-TIMS bulk baddeleyite Quadyuk Is. gabbro sill Heaman et al., 1992 Baffin Island 720* 4 4 Pb-Pb ID-TIMS bulk baddeleyite Cumberland dike dike Heaman et al., 1992 Coronation Gulf 720* 6 6 Pb-Pb ID-TIMS bulk baddeleyite Rae Group sill Heaman et al., 1992 Borden 720 8 8 U-Pb ID-TIMS bulk baddeleyite Borden dikes dike Pehrsson & Buchan, 1999 Greenland 721 4 4 U-Pb ID-TIMS 1 & 2 grain baddeleyite b.f. Qaanaaq dike dike Denyszyn et al., 2009b Ellesmere Island 721 2 2 U-Pb ID-TIMS 1 & 2 grain baddeleyite b.f. Cadogan dike dike Denyszyn et al., 2009b Victoria Island 721* 5 2 upper-intercept bulk baddeleyite Up. Sills, Shaler Gp sill Heaman et al., 1992 Greenland 726* 24 24 upper-intercept bulk baddeleyite Belcher Glacier dike dike Denyszyn et al., 2009b *Pb-Pb and upper intercept ages recalculated from isotope ratios using the U decay constant ratio of Schoene et al. (2006). b.f. = bulk fraction

Table S1: Geochronological constraints on the onset of the Sturtian glaciation and the Franklin Large Igneous Province (69–83). Table S2: Standard parameters used in the volcanic plume, climate and aerosol microphysics mod- eling. Parameter Values Plume entrainment coefficient α 0.1 Specific heat capacity of atmosphere at constant pressure cp,a 1040 J/K/kg Specific heat capacity of volatiles in plume cp,v 1900 J/K/kg Specific concentration of fines in eruption qf 0.01 kg/kg Specific concentration of volatiles in eruption qv 0.002 kg/kg 3 Magma density ρmagma 2900 kg/m Magma specific heat capacity cmagma 1100 J/K/kg −2 Solar flux FV 1284 W m Surface pressure ps 1.0 bar Ice-free global mean surface albedo A0 0.24 Ice-covered global mean surface albedo Ai 0.6 Total ice loss temperature T0 290 K Global ice coverage temperature Ti 260 K Aerosol droplet H SO specific concentration q 0.75 kg/kg 2 4 H2SO4 Sulfuric acid density ρ 1840 kg/m3 H2SO4 Water density ρ 1000 kg/m3 H2O H SO -air collision diameter d 4.3 × 10−10 m 2 4 H2SO4 −3 Plume condensation nuclei parameter nCn 100 particles cm Troposphere-stratosphere exchange timescale τexchange 1.0 y Mean stratosphere temperature Tstrat 218.6 K 3 Mean stratosphere pressure pstrat 5.3 × 10 Pa 5 3 Mean stratosphere OH concentrationn ˜OH 5.5 × 10 molecules/cm Mean stratosphere H O concentrationn ˜ 7.0 × 1012 molecules/cm3 2 H2O

Table S3: Chemical reaction rates used in the aerosol microphysical model. When used, M is assumed to be background air. The rate coefficient of reaction (1) was checked against the more general formula given in (84) and found to agree closely under the stratospheric conditions studied. The rate coefficient of reaction (4) was calculated by integrating loss rate values given in (56) over the altitude range of the box model. # Reaction Rate coefficient Reference −13 17 6 −1 1 SO2 + OH + M → HSO3 + M 8.2 × 10 /(7.0 × 10 + [M]) cm s (56) −11 3 −1 2 HSO3 + OH → SO3 + H2O 1 × 10 cm s (56) −15 3 −1 3 SO3 + H2O + H2O → H2SO4 + H2O 1.2 × 10 cm s (84, 85) −12 −1 4 H2SO4 + hν → SO2 + products 3.8 × 10 s (56)

17 Table S4: Results from the aerosol microphysical model when key parameters are varied. hMSi and hreff i are the time-mean total S aerosol burden and particle radius, respectively. All simulations are run for three years with yearly eruptions that inject 100 Mt of SO2 (5× Pinatubo). Simulation hMSi (Mt) hreff i (µm) Standard case, 5× Pinatubo 27.3 0.67 4 3 n˜OH = 5.5 × 10 molecules/cm 11.2 0.68 6 3 n˜OH = 5.5 × 10 molecules/cm 31.0 0.63 −3 nCn = 10 particles cm 24.3 0.83 −3 nCn = 1000 particles cm 28.3 0.63 τexchange = 0.5 y 17.1 0.57 τexchange = 2 y 35.8 0.75

18 717 Ma


I SA SC M EA Si NC NS L Franklin Large K ES Igneous Province G R

C Sf S A B Aw WA Ae

Figure S1: Palaeogeographic reconstruction for Rodinia at 717 Ma and distribution of the Franklin LIP. Modified from Li et al., (64). I = India; SA = southern Australia; NA = northern Australia; T = Tarim; SC = South China; M = Mongolia; EA = East Antarctica; Si = Siberia; NC = North China; L = Laurentia; NS = North Slope; ES = East ; G = Greenland; B = ; A = Amazonia; WA = West Africa; Aw = western Avalonia; Ae = eastern Avalonia; Sf = So Francisco; C = Congo; K = Kalahari; R = Rio Plata.

19 106


104 Sulfur (ppm)


102 -30 -20 -10 0 10 20 30

δ34S (‰ VCDT)

Diabase sills Calc- intrusions Sul de mineralization Carbonate Evaporites Natkusiak basalt Chilled margins Skarns Siliciclastic sedimentary rocks

Figure S2: Sulfur concentration and isotope data from the Natkusiak magmatic assemblage, sed- imentary rocks of the Shaler Supergroup, and contact metamorphic rocks. Data plotted from geochemical database of Franklin sills, Natkusiak Basalts and Shaler Supergroup rocks (15).

20 70


50 with O3 US STD US TROPIC 40

z [km] 30



0 180 200 220 240 260 280 300 T [K]

Figure S3: Radiative-convective equilibrium temperature profile for Earth’s atmosphere under ◦ cloud-free conditions with CO2 molar concentration of 300 ppmv, solar zenith angle of 60 , 1 bar atmospheric pressure and present-day insolation. The red line shows model results with ozone UV absorption included. Dashed and dotted lines show the US standard mid-latitude and tropical atmosphere, respectively.

21 70



40 z [km] 30



0 120 140 160 180 200 220 240 260 280 300 320 T [K]

Figure S4: Atmospheric temperature profiles used for the plume calculations displayed in Fig. 2 in the main text. In each case the climate model was run until radiative-convective equilibrium was reached.

22 1 1 plume 0.8 0.8 ambient

0.6 0.6

z [km] 0.4 z [km] 0.4

0.2 0.2

0 0 0 200 400 600 300 350 400 450 r [m] T [K]

1 1 plume 0.8 0.8 ambient

0.6 0.6

z [km] 0.4 z [km] 0.4

0.2 0.2

0 0 0 2 4 6 8 10 0 0.5 1 1.5 u [m/s] ρ [kg/m3]

Figure S5: (top left) Plume radius (top right) temperature (bottom left) plume vertical speed and (bottom right) density vs. altitude as simulated by the axisymmetric plume model. In the bottom right panel, the circle and asterisk indicate the maximum plume height and the neutral buoyancy 7 height, respectively. For this simulation we assumed Q0 = 5.6 × 10 J/s, r0 = 3 m, Tg,0 = 525 K and u0 = 10 m/s and the same thermodynamic parameters as (45).

23 30 T = 274 K s T = 280.6 K s 25 T = 290 K s T = 307.3 K s


present-day tropical tropopause


10 maximum plume height [km]


0 101 102 103 104 105 106 volume eruption rate [m3/s]

Figure S6: Same as Fig. 2 in the main text but assuming 10% of present-day ozone abundance.

24 30

T = 273 K s T = 279.2 K s 25 T = 287.6 K s

T = 301.7 K s 20

present-day tropical tropopause


10 maximum plume height [km]


0 101 102 103 104 105 106 volume eruption rate [m3/s]

Figure S7: Same as Figure S6 but also assuming background N2 pressure of 0.8 bar.

25 0.5


m] 0.3 µ [ 0.2 mean r


0 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 t [y]

12 aerosol 10 gas total 8

[Mt] 6 S M 4


0 0 0.1 0.2 0.3 0.4 0.5 0.6 0.7 0.8 0.9 1 t [y]

Figure S8: (top) Aerosol effective radius and (bottom) stratospheric sulfur loading as a function of time simulated by the aerosol microphysics model. The model initial conditions correspond to a single Pinatubo-like injection of 20 Mt SO2 into the stratosphere.

26 1.2


0.8 m] µ

[ 0.6 mean r 0.4


0 0 0.5 1 1.5 2 2.5 3 t [y]

102 [Mt] S 1 M 10

100 0 0.5 1 1.5 2 2.5 3 t [y]

Figure S9: (top) Aerosol effective radius and (bottom) stratospheric aerosol S burden vs. time for the standard set of 1×, 5× and 25× Pinatubo aerosol forcing simulations (corresponds to Fig. 3A in the main text).

27 0


-20 ] 2 -30

F [W/m -40 ∆ -50 ∆ F = -2.1 W/m2 avg ∆ F = -5.6 W/m2 -60 avg ∆ F = -7.6 W/m2 avg -70 0 0.5 1 1.5 2 2.5 3 t [y]




-30 ] 2 -40

-50 F [W/m

∆ -60

∆ F = -3.2 W/m2 -70 avg ∆ F = -8.6 W/m2 -80 avg ∆ F = -14 W/m2 avg -90 0 0.5 1 1.5 2 2.5 3 t [y]

Figure S10: Same as Fig. 3A in the main text, except for pre-aerosol tropical planetary albedo of 0.3 (top plot) and 0.2 (bottom plot). The lines in the plot correspond to tropical mean radiative forcing, while the values in the legend are global means.

28 0 ] 2 -20

-40 F [W/m 2 ∆ ∆ F = -2.4 W/m avg ∆ F = -5 W/m2 -60 avg ∆ F = -3.5 W/m2 avg -80 0 0.5 1 1.5 2 2.5 3 t [y]

0 ] 2 -20

-40 F [W/m 2

∆ ∆ F = -1.4 W/m avg ∆ F = -2.4 W/m2 -60 avg ∆ F = 1.1 W/m2 avg -80 0 0.5 1 1.5 2 2.5 3 t [y]

Figure S11: Same as Fig. 3A in the main text, except assuming a tropopause height of 10 km and simulation box domain from 30◦ N to 80◦ N, corresponding to the situation following a high latitude basaltic fissure eruption. The pre-aerosol mid-latitude planetary albedo is 0.365 (present-day value) in the top plot and 0.5 (high ice/snow cover case) in the bottom plot. At the highest eruption rate in the bottom plot, the combination of large aerosol particles and high pre-aerosol albedo leads to a small net warming.