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The evolution of volcanic systems following sector collapse Watt, Sebastian

DOI: 10.1016/j.jvolgeores.2019.05.012 License: Creative Commons: Attribution-NonCommercial-NoDerivs (CC BY-NC-ND)

Document Version Peer reviewed version Citation for published version (Harvard): Watt, S 2019, 'The evolution of volcanic systems following sector collapse', Journal of Volcanology and Geothermal Research, vol. 384, pp. 280-303. https://doi.org/10.1016/j.jvolgeores.2019.05.012

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Download date: 06. Oct. 2021 1 2 3 4 1 The evolution of volcanic systems following sector collapse 5 2 6 7 3 Sebastian F.L. Watt 8 4 School of Geography, Earth and Environmental Sciences, University of Birmingham, 9 10 5 Edgbaston, Birmingham B15 2TT, U.K. 11 12 6 e-mail: [email protected] 13 7 tel: 0121 414 6131 14 15 8 16 9 Abstract 17 18 10 Sector collapses affect volcanic edifices across all tectonic settings and involve a rapid 19 20 11 redistribution of mass, comparable in scale to the largest magmatic eruptions. The eruptive 21 12 behaviour of a volcano following sector collapse provides a test of theoretical relationships 22 23 13 between surface loading and magma storage, which imply that collapse-driven unloading 24 14 may lead to changes in eruption rate and erupted magma compositions. Large sector collapses 25 26 15 are infrequent events globally, with all historical examples being relatively small in 27 28 16 comparison to many of the events documented in the geological record. As a result, 29 17 exploration of the impacts of sector collapse on eruptive behaviour requires detailed 30 31 18 investigation of prehistoric collapses, but this is often hindered by poorly-resolved 32 19 stratigraphic relationships and dating uncertainties. Nevertheless, observations from a number 33 34 20 of volcanoes indicate sharp changes in activity following sector collapse. Here, a global 35 21 synthesis of studies from individual volcanoes, in both arc and intraplate settings, is used to 36 37 22 demonstrate a number of common processes in post-collapse volcanism. Multiple examples 38 3 39 23 from large (>5 km ) sector collapses in arc settings show that collapse may be followed by 40 24 compositionally anomalous, large-volume and often effusive eruptions, interpreted to 41 42 25 originate via disruption of a previously stable, upper-crustal reservoir. These anomalous 43 26 eruptions highlight that magma compositions erupted during periods of typical (i.e. 44 45 27 unperturbed by sector collapse) volcanism may not be representative of the range of 46 47 28 compositions stored within a vertically extensive crustal reservoir. If eruptible magma is not 48 29 present, upper-crustal reservoirs may rapidly solidify following collapse, without further 49 50 30 eruption, allowing more mafic compositions to ascend to the surface with only limited upper- 51 31 crustal modification, resulting in edifice regrowth at temporarily elevated eruption rates. 52 53 32 Subsequent re-establishment of an upper-crustal reservoir further supports a relationship 54 55 33 between surface loading and crustal storage, but long-term chemical and mineralogical 56 34 differences between pre- and post-collapse evolved magmas imply that a newly-developed 57 58 1 59 60 61 62 63 35 reservoir can overprint the influence of a preceding reservoir, forming a spatially and 64 36 compositionally distinct plumbing system. These broad patterns are replicated in intraplate 65 66 37 settings, despite differences in scale and melting processes; current evidence suggests that 67 38 post-collapse evolution of intraplate volcanoes can be explained by unloading-induced 68 69 39 destabilisation of the magma plumbing system, rather than increased melt production. What 70 71 40 emerges from an apparently diverse set of observations is a systematic behaviour that 72 41 strongly supports a coupling between edifice growth and magma ascent, storage and 73 74 42 pressurisation. Eruption rates, erupted compositions, and the style of volcanism at any 75 43 particular system may thus be modulated from the surface, and long-term shifts in surface 76 77 44 behaviour may thus occur without any changes in the deep parts of magmatic systems. 78 79 45 Observations of sharp post-collapse changes in erupted compositions, including the ascent of 80 46 primitive mafic magmas, also require a crystal-dominated mid- to upper-crustal reservoir, 81 82 47 consistent with recent models of crustal magmatic systems. 83 48 84 85 49 Keywords: sector collapse; magma storage; eruptive behaviour; edifice growth and 86 87 50 destruction; debris avalanche 88 51 89 90 52 Highlights 91 53 • Global synthesis of volcano-magmatic evolution following sector collapse 92 93 54 • Decompression-driven reservoir disruption leads to anomalous post-collapse 94 95 55 eruptions 96 56 • Rapid regrowth via mafic volcanism; subsequent re-establishment of shallow storage 97 98 57 • Implies direct modulation of shallow plumbing system development by surface load 99 100 58 • Common patterns in intraplate and arc settings 101 59 102 103 60 1. Introduction 104 61 Volcanic edifices across all tectonic settings are prone to structural instability and the 105 106 62 generation of large landslides (Ui, 1983; Siebert, 1984), resulting in a redistribution of 107 108 63 volcanic rock across the surrounding land surface. Landslides formed by edifice collapse 109 64 span a wide range of dimensions, and their scars have been identified on volcanoes ranging 110 111 65 from submerged seamounts to large ocean islands, and across subaerial composite volcanoes 112 66 in both arc and intraplate settings (Siebert et al., 1987, 2006; Moore et al., 1989; Deplus et al., 113 114 67 2001; Coombs et al., 2007; Staudigel and Clague, 2010). The triggers for such landslides are 115 116 117 2 118 119 120 121 122 68 varied. Although some, such as the sector collapse of Mount St Helen’s in 1980, are directly 123 69 associated with large magmatic eruptions (Glicken, 1996), structural failure is not always 124 125 70 linked with magma ascent (Siebert et al., 1987; McGuire, 1996). 126 71 127 128 72 Historical observations and deposit characteristics suggest that structural failure of volcanic 129 130 73 edifices generally occurs in a sudden, catastrophic event (although this may follow a long 131 74 period of more gradual flank spreading (e.g., Moore et al., 1989; Neri et al., 2004; Wooller et 132 133 75 al., 2004; Karstens et al., 2019); and failure itself may occur over several, shortly-spaced 134 76 stages (Glicken, 1996; Hunt et al., 2013)). The base of the failure plane in large edifice 135 136 77 collapses may lie deep within the volcano structure (Crandell, 1989; Glicken, 1996; Watt et 137 138 78 al., 2014) and even intersect basement rock (e.g., Wadge et al., 1995; Shea et al., 2008), and 139 79 in many cases cuts the central conduit. The mobilised mass may be remarkably large, in some 140 141 80 cases accounting for >10% of the edifice volume. Sector collapses thus profoundly alter both 142 81 the morphology of a volcano and the distribution of mass above an active magma plumbing 143 144 82 system, potentially reducing the thickness of overlying rock by a kilometre or more. 145 146 83 147 84 Theoretical analyses suggest that mass redistribution following edifice collapse can influence 148 149 85 pressurisation and failure conditions in stored magma bodies (Pinel and Jaupart, 2005; Pinel 150 86 and Albino, 2013). It is therefore plausible that major collapses may be followed by changes 151 152 87 in eruption rate or style, or in the composition of erupted magma. Anecdotal evidence from 153 88 several individual volcanoes supports this idea (e.g., Tibaldi, 2004; Hora et al., 2009; 154 155 89 Manconi et al., 2009), but it is not clear that there is a common pattern to post-collapse 156 157 90 activity. In addition, some volcanoes show no apparent change in behaviour following large- 158 91 scale edifice failure (e.g., Ponomareva et al., 2006; Zernack et al., 2012). Given the diversity 159 160 92 of volcanic systems affected by edifice collapse, this is perhaps unsurprising, but the limited 161 93 current understanding of the impacts of sector collapse on volcano-magmatic processes 162 163 94 provides the motivation for this work. This paper draws on existing data on volcanoes 164 165 95 affected by large-scale sector collapse to assess evidence for changes in subsequent volcanic 166 96 activity, how this varies between volcano-tectonic settings, and whether any common 167 168 97 processes or patterns of behaviour can be deduced. 169 98 170 171 99 2. Volcano sector collapse 172 173 100 The focus of this paper is on the impact that major volcanic structural failure has on the 174 101 associated magmatic system, rather than on the mechanisms of the failure itself or of the 175 176 3 177 178 179 180 181 102 mass-movement that it generates. Some brief terminological definitions are nevertheless 182 103 necessary, given the range of terms used in published literature to describe volcanic collapses 183 184 104 and their resultant deposits. 185 105 186 187 106 2.1. Terminology 188 189 107 All of the events documented in this paper involve the rapid gravity-driven transport of 190 108 material away from a volcanic edifice, in a process that is effectively instantaneous relative to 191 192 109 the timescale of volcano growth. The term sector collapse is used to describe the structural 193 110 failure itself. Early uses of this term were used principally to describe major structural 194 195 111 failures of composite arc volcanoes, where failed sectors of the cone generally encompassed 196 197 112 the central conduit and summit of the volcano (e.g., Ui, 1983; Siebert, 1984; Siebert et al., 198 113 1987), providing a useful distinction from smaller mass movements that affect a single flank 199 200 114 but don’t extend to the central vent, which may be termed flank collapses. The term lateral 201 115 collapse is also widely used, and emphasises the transport of mass away from the volcano, in 202 203 116 contrast with the subsidence that accompanies caldera formation. 204 205 117 206 118 In general, sector collapses on polygenetic volcanoes result from large (as a proportion of the 207 208 119 entire edifice) and deep-seated instabilities (Fig. 1), whereas some flank collapses are 209 120 relatively superficial landslides, and not of interest here due to their small dimensions. The 210 211 121 term sector collapse is thus used to emphasise that the studied events are large in the context 212 122 of the parent edifice, and a minimum primary volume criterion of 1 km 3 is applied here. This 213 214 123 value is somewhat arbitrary (and volumes of prehistoric events are often poorly constrained), 215 216 124 but this limits the events under consideration to those that are at least equivalent in volume to 217 125 a large magnitude (i.e. Magnitude >5; cf. Pyle, 2000) magmatic eruption. 218 219 126 220 127 Landslides on intraplate ocean islands form some of the largest mass-movements on Earth 221 222 128 (Moore et al., 1989; Masson et al., 2002), potentially exceeding hundreds of cubic kilometres 223 224 129 and dwarfing those on arc volcanoes (Fig. 1). The tectonic setting and magma generation 225 130 process at these long-lived volcanic systems leads to islands that are morphologically very 226 227 131 different to arc stratovolcanoes. Ocean islands typically have a lower overall gradient than arc 228 132 volcanoes, and may be dominated by rift zones (Carracedo, 1994; Walter et al., 2005), 229 230 133 potentially without a well-developed central magma system. Rift zone instabilities may be 231 232 134 deeply rooted and lead to large flank landslides (not strictly meeting the sector collapse 233 135 definition above), which nevertheless involve a smaller proportion of the total volcanic (i.e. 234 235 4 236 237 238 239 240 136 island) volume than many sector collapses in arc settings (Watt et al., 2014), despite their 241 137 extreme size (Fig. 2). Such mass movements clearly have the potential to form a significant 242 243 138 local stress perturbation, and are therefore considered here alongside their counterparts on 244 139 composite arc volcanoes. 245 246 140 247 248 141 The term debris avalanche is widely used to describe volcanic landslides (and their resultant 249 142 deposits) that involve non-juvenile material. However, the style of landslide movement can 250 251 143 vary widely during transport (Scott et al., 2001) (and may involve failure over several 252 144 discrete stages, particularly in ocean-island settings; cf. Hunt et al., 2013) and is also 253 254 145 dependent on the nature of the failed mass, resulting in a wide range of deposit morphologies 255 256 146 (Shea et al., 2008; Dufresne and Davies, 2009). The products of small volcanic landslides can 257 147 also form block-rich debris avalanches (e.g. Voight and Sousa, 1994; Voight et al., 2002; 258 259 148 Belousov et al., 2010), and at some volcanoes these may occur repeatedly on timescales as 260 149 short as a few hundred years (e.g. Begét and Kienle, 1992). The term thus has no overall 261 262 150 implications in terms of landslide volume or frequency. 263 264 151 265 152 266 267 153 2.2. Causal mechanisms 268 154 Ultimately, sector collapses are driven by gravitational instabilities arising from the 269 270 155 geologically rapid construction of relatively steep topography, but are also influenced by 271 156 internal structural and lithological heterogeneities (including basal discontinuities and 272 273 157 spreading; van Wyk de Vries et al., 2001; Shea et al., 2008), alteration (Reid et al., 2001) and 274 275 158 changes in pore-fluid pressure (Day, 1996), magmatic activity (Siebert et al., 1987), and 276 159 external destabilising processes such as the retreat of glacial ice (Watt et al., 2009a; Tormey, 277 278 160 2010) or, possibly, sea level changes (cf. McMurtry et al., 2004; Hunt et al., 2014; Coussens 279 161 et al., 2016; Paris et al., 2018). The precise trigger of sector collapses is generally difficult to 280 281 162 identify, and may not be attributable to a single process. However, many sector collapses are 282 283 163 not necessarily driven by magma ascent. Although some events, such as that at Mount St 284 164 Helens in 1980, are a direct consequence of shallow intrusive or eruptive processes 285 286 165 destabilising edifice flanks (Siebert et al., 1987), the role of eruptive activity in collapse is 287 166 ambiguous in many other cases, even when sector collapses occur during a period of eruption 288 289 167 (such as the 2018 collapse of Anak Krakatau, Indonesia). Other collapses show no evidence 290 291 168 of associated magmatic activity, and examples such as the 1888 collapse of Bandai, Japan, 292 169 which hasn’t had a magmatic eruption since 25 ka (Yamamoto et al., 1999), suggest that 293 294 5 295 296 297 298 299 170 sector collapses can occur entirely as a result of other destabilising processes. This is 300 171 significant, because it implies that patterns of eruptive behaviour spanning sector collapses 301 302 172 can provide a test of how magmatic systems respond to external stress changes, and how the 303 173 changes in surface loading that accompany edifice growth influence crustal magmatic 304 305 174 processes. For many prehistoric collapses, the role of magmatism in triggering sector collapse 306 307 175 cannot be determined. Although the presence of juvenile material mixed within collapse 308 176 deposits strongly suggests an eruption-driven process, the same is not true of juvenile 309 310 177 products that immediately overlie collapse deposits, since they could reflect collapse- 311 178 triggered eruption rather than the reverse. Such ambiguities in driving mechanisms and the 312 313 179 contemporaneous state of the underlying magma reservoir present a challenge in evaluating 314 315 180 the general effects of sector collapses on crustal magmatic processes. 316 181 317 318 182 2.3. Global distribution 319 183 Many individual sector collapses, and in particular their deposits, have been studied in detail 320 321 184 (e.g. Wadge et al.,1995; Glicken, 1996; Shea et al., 2008; Hora et al., 2009). The majority of 322 323 185 these well-studied examples are from subaerial arc volcanoes, although a significant number 324 186 have also been documented from intraplate ocean islands such as Hawaii and the Canary 325 326 187 Islands (Moore et al., 1989; Masson et al., 2002). Individual events are commonly recognised 327 188 both from their hummocky deposits, which transport poorly-sorted and block-rich material 328 329 189 tens of kilometres away from the volcano, and their characteristic amphitheatre-shaped scars 330 190 (Siebert et al., 1987). The large number of identified examples indicates that collapse is a 331 332 191 ubiquitous process affecting constructive volcanic landforms. 333 334 192 335 193 A global compilation of documented sector collapses, shown in Fig. 2, comprises over 300 336 337 194 events with estimated volumes >1 km 3, from over 200 individual volcanoes. Event volumes 338 195 are often very poorly constrained, due to infilling or erosion of scars and limited information 339 340 196 on deposit thicknesses. For events where specific volume estimates are available (Fig. 2), 341 342 197 71% are from subduction zone settings, 25% from intraplate ocean islands, and 4% from 343 198 other settings. Compared to the global distribution of subaerial volcanoes (Global Volcanism 344 345 199 Program, 2013), these values suggest that intraplate ocean-island landslides are over- 346 200 represented in the dataset, since they account for 7% of volcanoes. Rift and continental 347 348 201 intraplate volcanoes are under-represented, and there is also a slight under-representation of 349 350 202 events in island-arc settings. These patterns are likely to reflect a sampling bias: the few 351 203 regional-scale surveys of island arcs (Deplus et al., 2001; Coombs et al.. 2007; Silver et al., 352 353 6 354 355 356 357 358 204 2009) show that sector collapse deposits occur around the majority of volcanic islands, but 359 205 their identification is dependent on good quality bathymetric data. The over-representation of 360 361 206 ocean-island collapses partly reflects the large number of events documented from Hawaii 362 207 (Moore et al., 1989), the Canary Islands (Masson et al., 2002; 2006), the Cape Verde Islands 363 364 208 (Masson et al., 2008) and Réunion (Oehler et al., 2008). It also reflects the long lifespan of 365 366 209 many ocean islands and the large dimensions of their landslides, meaning that relatively 367 210 ancient deposits can still be identified while those of a similar age in subduction-zone settings 368 369 211 are more likely to have been buried, or their scars eroded. 370 212 371 372 213 In subduction zone settings, over 50% of documented sector collapses (only considering 373 3 374 214 those >1 km and with a specified volume) have volumes between one and five cubic 375 215 kilometres, with a thick tail of larger events that extends up to 50 km 3 (Fig. 2). This pattern 376 377 216 mirrors the overall size distribution of typical arc volcanoes: a compilation of 400 subaerial 378 217 arc stratovolcanoes (Grosse et al., 2014) indicates a median edifice volume of 20.1 km 3 but a 379 380 218 mean volume of 43 km 3, skewed by a small number of very large edifices that exceed 200

381 3 382 219 km in volume. There are several examples of sector collapses that mobilise 10% of the total 383 220 edifice volume, with some affecting over 25% of the total volcanic structure (e.g. Wadge et 384 5-6 385 221 al., 1995; Zernack et al., 2012). Given typical lifetimes of 10 years for individual 386 222 subduction zone volcanoes (White et al., 2006), a single event may transport a volume 387 388 223 equivalent to tens of thousands of years of accumulated material. Such collapses are thus 389 224 major and relatively infrequent events in the history of individual volcanoes. 390 391 225 392 393 226 3. General evidence of magmatic responses to external stress changes 394 227 Several external processes have been proposed as causes of significant stress perturbations to 395 396 228 stored magma. The best studied of these is ice retreat, particularly following the end of the 397 229 last glaciation. Evidence from Iceland suggests a substantial increase in volcanism in the late 398 399 230 Pleistocene and early Holocene, attributed principally to enhanced melt production (Jull and 400 401 231 McKenzie, 1996; Maclennan et al., 2002; Eason et al., 2015). 402 232 403 404 233 Glacial unloading is different from sector-collapse induced unloading in two ways. Firstly, 405 234 deglaciation is a regional-scale effect, and the resultant lithospheric adjustment may drive 406 407 235 increased mantle melting (in a rift setting such as Iceland, at least; evidence is more limited in 408 409 236 arc settings (cf. Watt et al., 2013; Rawson et al., 2016)), as well as affecting crustally stored 410 237 magma. In contrast, volcanic edifices represent a local perturbation to the lithostatic stress 411 412 7 413 414 415 416 417 238 field (Pinel and Jaupart, 2005). The impact of sector collapse is thus limited to upper crustal 418 239 magma bodies in arc settings, and is not likely to affect melting processes, with the possible 419 420 240 exception of large collapses in ocean-island settings (Presley et al., 1997; Hildenbrand et al., 421 241 2004; Manconi et al., 2009) (see Section 8, below). Secondly, the timescale of sector collapse 422 423 242 is much more rapid than ice retreat, and the stress change thus exceeds the rate of magmatic 424 425 243 processes. This means that sector collapse provides a much clearer natural laboratory than ice 426 244 retreat to test the impact of surface unloading on stored magma. The effectively instantaneous 427 428 245 nature of collapse means that any effects can be placed in a clear stratigraphic context, even if 429 246 absolute ages are unknown, providing a sharp distinction between pre- and post-unloading 430 431 247 conditions. 432 433 248 434 249 The rapid nature of sector-collapse induced stress changes is comparable to earthquake 435 436 250 induced magmatic stresses (Albino, 2011), but the local static stress change is far larger. A 437 251 range of evidence supports magma disruption following earthquakes, leading to triggered 438 439 252 eruptions within days (Linde and Sacks, 1998; Manga and Brodsky, 2006), to potentially 440 441 253 years (Watt et al., 2009b). However, most postulated mechanisms suggest that these effects 442 254 are due to dynamic stress induced by seismic waves, rather than static stress changes (Hill et 443 444 255 al., 2002; Walter and Amelung, 2007). 445 256 446 447 257 4. Theoretical effects of sector collapse on magmatic systems 448 258 The relationship between magma storage, pressurisation and local crustal stress has been 449 450 259 explored through a variety of analytical and numerical approaches (Pinel and Jaupart, 2003, 451 452 260 2005; Gudmundsson, 2006; Manconi et al., 2009; Karlstrom et al., 2015). These models 453 261 generally assume simplified geometries and physical properties of magma and surrounding 454 455 262 rock. Such simplification is justifiable given that constraints on subsurface storage conditions 456 263 are limited by the spatial resolution and uncertainties of geophysical interpretations (e.g., 457 458 264 Foroozan et al., 2010; Paulatto et al., 2012). Petrological studies are increasingly highlighting 459 460 265 a complex picture of magma transport, storage and crystallisation (Cashman and Blundy, 461 266 2013), suggesting that magma reservoirs may depart significantly from the simplified liquid 462 463 267 chamber of theoretical models (Cashman et al., 2017). Nevertheless, it is clear that eruptions 464 268 are fed by large and relatively homogeneous volumes of eruptible, melt-dominated magma, 465 466 269 regardless of how and when these are assembled, and theoretical approaches can thus provide 467 468 270 valuable insights into how the volcanic edifice load, or surface loading in general, influences 469 271 storage conditions and dyke formation. 470 471 8 472 473 474 475 476 272 477 273 4.1. Loading, magma storage and dyke formation 478 479 274 An edifice load imposes a departure from a lithostatic stress state in the underlying crust. This 480 275 effect decreases with depth and becomes negligible at around three times the edifice radius (if 481 482 276 the edifice is supported by the strength of the lithosphere; i.e. the edifice radius is small 483 484 277 relative to lithospheric elastic thickness) (Pinel and Jaupart, 2005), and is thus most relevant 485 278 to upper-crustal magma bodies in arc settings. In affecting the local stress field, the edifice 486 487 279 load may influence magma chamber growth (Karlstrom et al., 2010; Gudmundsson, 2012), 488 280 dyke formation (Pinel and Jaupart, 2005; Hurwitz et al., 2009), and dyke propagation (Pinel 489 490 281 and Jaupart, 2000; Muller et al., 2001; Kervyn et al., 2009), and focus dyke ascent around a 491 492 282 central vent. Dyke formation – the critical precursor to eruption – may be modelled in terms 493 283 of a rupture criterion for a liquid body in an elastic medium (Fig. 3) (Pinel et al., 2010). 494 495 284 Rupture will occur at the maximum pressure above the lithostatic state (i.e. overpressure), Pr, 496 285 that the host rock can sustain. A linked variable, Pc, describes the pressure at which a dyke 497 498 286 will close and an eruption will cease (Pinel et al., 2010). Pr is dependent on reservoir 499 500 287 geometry and the stress state of the wall rock. Some models (Pinel et al., 2003;

501 288 Gudmundsson, 2006) relate Pr to the host-rock tensile strength (cf. Gudmundsson, 2012), but 502 503 289 others (Grosfils, 2007; Hurwitz et al., 2009; Gerbault et al., 2012) argue that much higher 504 290 overpressures can be sustained if gravitational loading is accounted for. 505 506 291 507 292 4.2. Edifice growth and eruption rate 508 509 293 Following the analytical approach of Pinel and Jaupart (2003), as an edifice increases in size 510 511 294 the load-related tensile stress on a subsurface liquid body increases to a maximum (and Pr 512 295 and Pc decrease to a minimum; Fig. 3). Beyond this point, increasing edifice growth 513 514 296 progressively hinders rupture. A broad range of factors, including reservoir shape, affect the 515 297 details of this relationship but not the qualitative principles. If an open link to a deeper 516 517 298 reservoir is assumed, then the rate of pressurisation of the upper reservoir will be determined 518 519 299 by the pressure difference between them (Fig. 3A). When Pr is relatively low, then following 520 300 an eruption, replenishment of the upper reservoir (i.e. the repressurisation from Pc to Pr; 521 522 301 assuming no replenishment during the eruption) will be relatively rapid, and the time to the 523 302 next eruption correspondingly short (i.e. a high theoretical eruption rate). This therefore 524 525 303 predicts that edifice growth is accompanied by an increasing eruption rate, to the point where 526 527 304 Pr reaches a minimum, and then a decreasing eruption rate (Pinel et al., 2010) (Fig. 3B). 528 305 There are large simplifications in this model, including assumptions of constant magma 529 530 9 531 532 533 534 535 306 properties and supply rates, and that dykes consistently feed eruptions rather than intrusions, 536 307 but it nevertheless provides a simple framework to consider the impacts of edifice growth and 537 538 308 destruction. It also implies a limit to edifice growth, with the surface load modulating the 539 309 storage system, which fits with broad observations of a finite lifetime to individual volcanoes 540 541 310 and the common occurrence of successively younger edifices adjacent to older, extinct 542 543 311 systems (e.g. Singer et al., 1997; Davidson and de Silva, 2000). 544 312 545 546 313 4.3. The impact of sector collapse 547 314 The simplified theoretical relationship between edifice size, dyke formation and eruption rate 548 549 315 allows a range of different scenarios to be postulated following the sudden reduction in 550 551 316 edifice load that accompanies edifice collapse (or more accurately, the redistribution of this 552 317 load across a broader area; Albino et al., 2010; Pinel and Albino, 2013) (Table 1). If sector 553 554 318 collapse is triggered by a shallow intrusion or incipient eruption (i.e. a Mount St. Helens type 555 319 event), then this implies that eruptible magma is present and that the initial Pr criterion has 556 557 320 been met. Collapse will always reduce the magma chamber pressure, Pm (Fig. 3), and the 558 559 321 nature of the subsequent eruption is dependent on the difference between the resultant Pr and

560 322 the post-collapse Pm. In some circumstances, a larger eruption may result, but Pinel and 561 562 323 Albino (2013) show that for many edifice and chamber geometries, the eruption will be 563 324 smaller or may stall entirely. For subsequent eruptions, the impact of collapse depends on 564 565 325 where the system lies on the Pr curve (Fig. 3) (Pinel and Albino, 2013). Broadly speaking, for 566 326 edifices at a later stage in the growth cycle outlined above, a load decrease will reduce P , 567 r 568 327 favouring conditions for dyke rupture and an increased subsequent eruption rate (Fig. 3). The 569 570 328 opposite effect is predicted for smaller edifices. Similar outcomes are predicted if sector 571 329 collapse is not intrusion-related, but if an eruptible body of magma is present. The newly 572 573 330 established stress state may, in some circumstances, result in a Pr that is less than the post- 574 331 collapse P , and in this instance an eruption would be triggered by the collapse. Finally, if no 575 m 576 332 eruptible magma is present, the post-collapse stress state will simply alter the conditions for 577 578 333 subsequent dyke formation, if and when an eruptible liquid body forms in the upper crust. 579 334 580 581 335 In addition to these effects, and regardless of the change in Pr, the reduction in Pm that 582 336 accompanies collapse will induce magma ascent from a deeper reservoir, if an open 583 584 337 connection exists (Pinel et al., 2013). Sector collapse may therefore initiate the addition of 585 586 338 heat or volatiles to the upper reservoir, which may potentially (but not necessarily for larger- 587 588 589 10 590 591 592 593 594 339 volume systems; cf. Ruprecht and Wörner, 2007)) lead to eruption or compositional changes 595 340 in subsequently erupted magmas. 596 597 341 598 342 In general, given that surface loading inhibits the ascent of denser mafic magmas and 599 600 343 promotes their stalling in the upper crust (Pinel and Jaupart, 2004), a plausible outcome of 601 602 344 collapse is the eruption of more mafic magmas, or of magmas that show a greater influence 603 345 of a deeper, mafic input. The principle of edifice and chamber co-development acting to 604 605 346 influence mafic dyke ascent is supported by observations of outlying vents at Mount 606 347 Mazama, which become increasingly restricted to locations distant from the central vent prior 607 608 348 to the climactic Crater Lake eruption (Karlstrom et al., 2015). In this case, the authors suggest 609 610 349 that development of the shallow magma system was driven by increased deep magma influx, 611 350 but the observation nevertheless supports a relationship between edifice and magma reservoir 612 613 351 growth, dyke capture, and a modulation of erupted magma compositions at different stages of 614 352 volcano development (cf. Schindlbeck et al., 2014). 615 616 353 617 618 354 The theoretical model outlined above (Fig. 3; Table 1) implies that the manifestation of sector 619 355 collapse on subsequent activity may not follow a common pattern, even before parameters 620 621 356 such as magma composition, density and volatile content are accounted for. If these are also 622 357 considered, it is clear that magmatic responses to sector collapse could be diverse, being 623 624 358 dependent on the local load-induced stress state (i.e. the relationship between edifice size and 625 359 magma reservoir geometry), the nature of the upper crustal magma reservoir at the time of 626 627 360 collapse, and the geometry and composition of the crustal plumbing system. Nevertheless, the 628 629 361 hypothesis that emerges is that sector collapse results in a sudden shift in eruptive behaviour 630 362 relative to pre-collapse activity, potentially manifested as a change in eruption frequency, 631 632 363 magnitude, composition or style. This would imply that changes in volcanic composition or 633 364 output do not necessarily reflect changes at the source, but can be modulated from the 634 635 365 surface. Past events can be examined in this framework. Conversely, the type of response, if 636 637 366 one can be identified, may provide constraints on the nature of the underlying magma 638 367 reservoir. 639 640 368 641 369 5. Volcano growth and eruption-rate measurements 642 643 370 The model outlined in Section 4 implies that changes in eruption rate may be a key 644 645 371 manifestation of sector collapse. However, eruption rates can only be accurately defined from 646 372 recent historical records; even at the best studied volcanoes, reconstructions of prehistoric 647 648 11 649 650 651 652 653 373 eruptions are still likely to be highly incomplete (Brown et al., 2014). Beyond historical 654 374 timescales, the mean volumetric eruptive flux calculated across a defined time period can be 655 656 375 used as a proxy for eruption rate. This measurement is reliant on detailed stratigraphic and 657 376 chronological information and is not a true indicator of eruption rate (since it reflects the 658 659 377 magnitude as well as the frequency of eruptions). It also fails to account for intrusive growth 660 661 378 of volcanic edifices. Nevertheless, the theoretical relationship in Fig. 3 implies that eruptive 662 379 flux correlates with eruption rate, and can thus be used to explore variations in eruptive 663 664 380 behaviour over prehistoric timescales. Before doing so, it is instructive to assess how well 665 381 eruptive flux can be determined on timescales of 10 3-4 years, both in order to interpret 666 667 382 observations at volcanoes affected by sector collapse and as a general test of the theoretical 668 669 383 relationship between edifice growth and eruption rate. 670 384 671 0-2 672 385 Measurements of eruptive flux are necessarily time-averaged. On short timescales (10 673 386 years), very high values can be maintained (Wadge, 1982), but there is a steady decline in the 674 675 387 estimated eruptive flux as the duration of the measurement increases (Fig. 4, inset). Thus, 676 677 388 estimates of eruptive flux made over the tens of thousands of years will hide finer-scale 678 389 variations, and particularly short episodes of elevated output (Hildreth and Lanphere, 1994) 679 680 390 such as those that may be expected following sector collapse. 681 391 682 683 392 There are relatively few studies that have attempted a detailed quantification of eruptive flux 684 393 over the lifetime of a volcano. Fig. 4 shows a compilation of thirteen such datasets, all from 685 686 394 composite arc volcanoes. Even for these comprehensive field studies, temporal resolution is 687 4 688 395 rarely <10 years. Erosion and burial, along with dating limitations, mean that age-volume 689 396 relationships are obscured by multiple uncertainties, and estimates of eruptive flux are likely 690 691 397 to be underestimated, with dispersed pyroclastic deposits difficult to account for. 692 398 Nevertheless, Fig. 4 shows that episodic growth behaviour is a common characteristic of 693 694 399 composite arc volcanoes, with relatively short periods characterised by volumetric output 695 696 400 rates as much as on order of magnitude greater than relatively quiescent interludes (cf. 697 401 Davidson and de Silva, 2000). This episodic pattern implies that magma systems remain 698 699 402 active across long timescales, even if little extrusion occurs (Hildreth and Lanphere, 1994). 700 403 Fig. 4 also highlights the decrease in data resolution with time: the highest measured values 701 702 404 occur within the past few ten thousand years; before this time, high-flux episodes are 703 704 405 generally not resolvable. The average eruptive flux across all these datasets shows that this is 705 406 not just a smoothing effect, because the long-term mean decreases markedly beyond ~120 ka. 706 707 12 708 709 710 711 712 407 This suggests that erosion, burial and dating limitations all hinder our ability to reconstruct 713 408 volcanism beyond the most recent episodes of cone-building (cf. Singer et al., 2008). 714 715 409 Reconstructions prior to the youngest period of volcanism at any volcano are thus likely to 716 410 significantly underestimate eruptive flux and are unlikely to resolve variability on timescales 717 718 411 below 10 4 years. The temporal resolution at which past activity can be reconstructed therefore 719 720 412 remains too coarse to fully test whether edifice growth cycles replicate in detail the shape of 721 413 the curve predicted in Fig. 3. 722 723 414 724 415 Although it is not always clear what divides individual constructional episodes, several 725 726 416 studies identify major sector collapse deposits that delineate periods of cone-building (Hall et 727 728 417 al., 1999; Thouret et al., 2005; Hora et al., 2007; Hall and Mothes, 2008; Samaniego et al., 729 418 2012) and cycles of collapse and regrowth characterise the history of many individual 730 731 419 volcanoes (Robin et al., 1990; Zernack et al., 2012; de Silva and Lindsay, 2015). Sector 732 420 collapse thus appears to be a fundamental process in composite volcano development, and 733 734 421 systematic variation in volcanic output may be related both to the maturation of a magma 735 736 422 reservoir and the growth of the overlying edifice, even if the details of how post-collapse 737 423 volcanism differs from pre-collapse activity are not documented. In many cases, collapse 738 739 424 scars are buried by rocks erupted during subsequent rapid regrowth (de Silva and Lindsay, 740 425 2015), providing general support for enhanced post-collapse activity. However, a more 741 742 426 detailed evaluation of this process requires reconstructions that can address whether, by 743 427 effectively shifting the coupled volcano-magma system to an earlier stage in a theoretical 744 745 428 growth cycle, collapse drove a resultant change in eruptive behaviour. To reject the null 746 747 429 hypothesis, we must ideally demonstrate that through the rebuilding process, a volcanic 748 430 system moves back towards its pre-collapse state, thus responding both to the perturbation 749 750 431 and to subsequent regrowth. 751 432 752 753 433 6. Historical events 754 755 434 Wide recognition of sector-collapse processes came about following the Mount St. Helens 756 435 eruption of 1980, although earlier eruptions such as those at Bandai (1888) and Bezymianny 757 758 436 (1956) had been identified as distinctive in terms of their deposits and horseshoe-shaped scars 759 437 (cf. Siebert, 1984). The largest historical sector collapse, at Ritter Island (1888), was 760 761 438 originally interpreted as a caldera-forming explosive eruption and only later recognised as a 762 763 439 sector collapse (Johnson, 1987). These historical events provide only a geologically short 764 440 window into subsequent activity, potentially insufficient to fully examine the consequences 765 766 13 767 768 769 770 771 441 of collapse on the magmatic system, but at least providing information at a high temporal 772 442 resolution and overcoming the age uncertainties and incompleteness that hinders prehistoric 773 774 443 reconstructions. 775 444 776 777 445 Historical collapses that meet a >1 km 3 volume criterion are shown in Table 2. Smaller 778 3 779 446 collapses, such as that at Bezymianny (0.5 km ; Belousov et al., 2007) or the December 2018 780 447 collapse of Anak Krakatau, are excluded from the table, but described below in the context of 781 782 448 other historical events. Recently revised volumes of the Bandai collapse (Yoshida, 2013) 783 449 suggest its volume was also <1 km 3, and the Oshima-Oshima (1741) collapse volume is 784 785 450 relatively poorly constrained (Table 2). Although these were all major landslides and caused 786 787 451 extensive local impacts, they are relatively small in the context of documented prehistoric 788 452 sector collapses (Fig. 2). 789 790 453 791 454 6.1. Eruption-associated collapses 792 793 455 With the exception of Bandai and possibly Ritter, all historical examples of sector collapse 794 795 456 were associated with fresh magma ascent. Collapse occurred during an explosive eruption at 796 457 Oshima-Oshima and Bezymianny, and during a longer-lasting phase of eruptive activity at 797 798 458 Anak Krakatau, while the Mount St. Helens collapse was preceded by magma ascent to a 799 459 shallow level within the edifice. Decompression of the shallow Mount St. Helens intrusion, 800 801 460 initiated by landslide movement, resulted in a powerful lateral blast (Glicken, 1996), and a 802 461 similar effect occurred following conduit depressurisation at Bezymianny (Belousov et al., 803 804 462 2007). At Shiveluch (and also at Harimkotan, 1933 (0.4 km 3; Belousov and Belousova, 805 806 463 1996)) magma was not sufficiently shallow for a directed blast to result (Belousov, 1995), but 807 464 collapse was preceded by accelerating seismicity and immediately followed by a large 808 809 465 explosive eruption, suggesting that an incipient eruption caused these collapses. 810 466 811 812 467 It is difficult to assess whether these historical eruptions proceeded differently relative to a 813 814 468 hypothetical scenario without collapse (cf. Table 1). However, it is clear that there are no 815 469 instances among them of a stalled eruption (assuming that the Bandai collapse was not due to 816 817 470 magma ascent, which is reasonable given an absence of magma extrusion since 25 ka), which 818 471 is one of the outcomes of collapse put forward by Pinel and Albino (2013) (Table 1). In all 819 820 472 the above cases, the intense explosive eruption that accompanied collapse was followed by an 821 822 473 effusive phase (and a phreatomagmatic phase in the case of Anak Krakatau) that partially 823 824 825 14 826 827 828 829 830 474 infilled the collapse scar, which appears to be a common process in eruption-associated sector 831 475 collapses (Table 2). 832 833 476 834 477 At Oshima-Oshima and Bezymianny, collapse followed a long period of quiescence (Satake 835 836 478 and Kato, 2001; Belousov et al., 2007; Girina, 2013). However, in both cases it was a 837 838 479 renewed eruptive phase that triggered collapse, rather than the reverse. At Oshima-Oshima 839 480 (and similarly at Harimkotan; Belousov and Belousova, 1996), this eruptive phase declined 840 841 481 over the subsequent fifty years (Katsui and Yamamoto, 1981), and there have been no 842 482 subsequent eruptions. In this case, there is nothing to indicate that the post-collapse state of 843 844 483 the system differs significantly from pre-collapse conditions. At Bezymianny, the initial 845 846 484 event heralded a period of intense activity that is ongoing today and has largely buried the 847 485 collapse scar (Girina, 2013), but since this renewed phase of volcanism began prior to the 848 849 486 collapse, its origin is likely related to deeper magmatic processes. 850 487 851 852 488 6.2. The largest historical collapses: Ritter, Shiveluch and Mount St. Helens 853 854 489 In contrast to Oshima-Oshima and Bezymianny, Shiveluch, Ritter and Mount St. Helens were 855 490 relatively active volcanoes before collapse and remain so today (Global Volcanism Program, 856 857 491 2013; Day et al., 2015; Watt et al., 2019). At Shiveluch, the 1964 event is just one of multiple 858 492 Holocene sector collapses (Belousov et al., 1999; Ponomareva et al., 2006) and fits within a 859 860 493 long-term pattern of repetitive collapse and regrowth. Similar behaviour is suggested by 861 494 multiple relatively small (generally ≤1 km3) collapses at volcanoes such as Stromboli 862 863 495 (Tibaldi, 2001), Augustine (Begét and Kienle, 1992) and Harimkotan (Belousov and 864 865 496 Belousova, 1996)). Shiveluch is characterised by an unusually high magma flux (Belousov et 866 497 al., 1999), and the 1964 collapse volume is equivalent to ~10 2 years of magmatic output. In 867 868 498 this context, the collapse is relatively minor. It accounted for <1% of the total edifice volume 869 499 (cf. Grosse et al., 2014), and predominantly involved material extruded since the preceding 870 871 500 edifice failure. Regular and relatively small collapses, involving recently erupted material, 872 873 501 may be a common outcome of rapid construction at highly active volcanoes. The 1964 event 874 502 was not unusual in the context of Holocene activity at Shiveluch, and would not, therefore, be 875 876 503 expected to mark a major change in behaviour relative to pre-1964 activity. It is also dwarfed 877 504 by a late-Pleistocene or early-Holocene collapse, which was not accompanied by a major 878 879 505 eruption (Ponomareva et al., 2006) but does mark a distinct shift in erupted magma 880 881 506 compositions (Belousov et al., 1999; see Section 7). 882 507 883 884 15 885 886 887 888 889 508 As suggested at Shiveluch, the significance of sector collapse at a volcano may not be 890 509 correlated closely with collapse volume, but may be better evaluated in terms of long-term 891 892 510 eruptive flux and total edifice volume. In this sense, the Ritter and Mount St. Helens events 893 511 stand out among historical collapses, in mobilising 10–20% of the edifice (based on edifice 894 895 512 volume estimates from Grosse et al. (2014) and Day et al. (2015)). At Ritter, recent 896 897 513 observations suggest that the collapse marked a shift in erupted compositions and was also 898 514 immediately followed by a compositionally anomalous felsic explosive eruption (Watt et al., 899 900 515 2019). This post-collapse eruption was bimodal, containing a mafic phase alongside felsic 901 516 material that is distinctive in both its glass chemistry and mineral content (containing 902 903 517 phenocryst amphibole), and unlike any other known eruption products from Ritter. The 904 905 518 eruption is consistent with collapse-driven perturbation of the underlying magma reservoir 906 519 (see Section 7). Scoriaceous deposits from subsequent eruptions are less compositionally 907 908 520 diverse, but comprise a mixture of mafic clasts that are more primitive than pre-collapse 909 521 material, as well as intermediate compositions, suggesting tapping of discrete crustal magma 910 911 522 bodies during post-collapse regrowth of the volcano (Watt et al., 2019). It is unclear if these 912 913 523 changes were accompanied by a change in eruption rate; the volcano experienced frequent 914 524 minor basaltic explosive eruptions before 1888, and similar activity is ongoing within the 915 916 525 collapse scar (Day et al., 2015; Watt et al., 2019), implying no major changes in eruption 917 526 style. 918 919 527 920 528 Mount St. Helens is better studied than Ritter, and has been highly active throughout the 921 922 529 Holocene. The 1980 collapse followed a pause in activity of ~120 years, and principally 923 924 530 involved material erupted over the preceding 2.5 kyr, ranging from basalts to dacites 925 531 (Glicken, 1996). The collapse and explosive eruption was followed by extrusion of a highly 926 927 532 crystalline and gas-poor dacitic lava dome, with no major changes in magma chemistry, 928 533 suggesting that the event tapped a relatively well-mixed and homogeneous magma body. 929 930 534 Broadly similar dacitic magma was erupted in a subsequent dome-building phase from 2004- 931 932 535 2008 (Pallister et al., 2008), and both magmas are comparable to those erupted in pre- 933 536 collapse historical eruptions. The 2004-08 lavas show the least evidence of mixing with a 934 935 537 basaltic contaminant for any eruption in the past 500 years (Pallister et al., 2008), and there is 936 538 no strong evidence to support the input of fresh magma from the lower crust. 937 938 539 939 940 540 Recent eruptions at Mount St. Helens suggest a long-lived dacitic upper-crustal reservoir, 941 541 with no significant change in post-collapse eruption rates or magma compositions. One 942 943 16 944 945 946 947 948 542 notable observation, however, is that the equilibration pressure of the 2004 magma (130 949 543 MPa) is significantly lower than that of the magma that initiated the 1980 collapse (220 MPa) 950 951 544 (Pallister et al., 2008; Rutherford and Devine, 2008). Pallister et al. (2008) interpret this in a 952 545 framework of a magma reservoir extending from depths of 5 to 12 km, tapped at different 953 954 546 levels by different eruptions. The change in equilibration depth between successive eruptions, 955 956 547 in a shallowing pattern, was noted by Gardner et al. (1995) and has characterised two cycles 957 548 of activity at Mount St. Helens over the past 4000 years. It is, however, unclear why the 958 959 549 equilibration depths of magma batches follows a systematic shallowing trend, and Pinel and 960 550 Jaupart (2003; see also Pinel et al., 2010) offer an alternative explanation. They argue that 961 962 551 edifice destruction following large explosive eruptions or sector collapse results in a 963 964 552 reduction of the overpressure required to initiate dyke ascent, and that conversely, edifice 965 553 growth results in higher overpressures. In this model, the equilibration pressures do not 966 967 554 reflect depth, but the overpressure sustained in a fixed magma reservoir. Although this is 968 555 qualitatively consistent with theoretical relationships between edifice growth, chamber 969 970 556 pressurisation and replenishment rates (Pinel and Jaupart, 2003), the absolute magnitude of 971 972 557 the pressure range during eruptive cycles at Mount St. Helens (up to 180 MPa) is an order of 973 558 magnitude larger than most estimates of maximum magma reservoir overpressures 974 975 559 (Gudmundsson, 2012). Constraints from mineral assemblages and other petrological 976 560 indicators also support a vertically extensive reservoir, with magma phenocryst mixtures 977 978 561 spanning this pressure range (Rutherford and Devine, 2008; Cashman and Blundy, 2013). 979 562 The post-collapse behaviour of Mount St. Helens doesn’t, therefore, fit a simple model of 980 981 563 collapse-modulated pressurisation. However, it is possible that the reduction in equilibration 982 983 564 pressures is linked to the reduced edifice load, if loading influences the depth range of 984 565 magma convection, the accumulation depths of mobile magmatic components, or the stalling 985 986 566 depth of any mafic, replenishing magma. 987 567 988 989 568 7. Magmatic impacts of sector collapse in subduction zone settings 990 991 569 Many of the factors controlling arc magma compositions, including slab input, mantle 992 570 melting and lower crustal interactions, are not plausibly influenced by sector collapse (see 993 994 571 Section 4). Any collapse-associated changes in eruptive behaviour are thus expected to reflect 995 572 mid- to upper-crustal magma processing, including the timescale of crustal transit, storage 996 997 573 and assimilation, the relative proportions and influence of mixing, the ascent and stalling of 998 999 574 mafic magmas, and consequent pressure and temperature controls on mineral phase stability. 1000 575 Such processes may be apparent via changes in erupted compositions or eruption rate and 1001 1002 17 1003 1004 1005 1006 1007 576 style. Evaluating evidence of this for prehistoric collapses requires comparison between pre- 1008 577 and post-collapse magmas at a range of timescales, which is often hindered by limited 1009 1010 578 exposures and poorly resolved age relationships. Although many volcanoes show evidence 1011 579 for rapid post-collapse regrowth and scar burial (e.g., Hora et al., 2007; Samper et al., 2007; 1012 1013 580 de Silva and Lindsay, 2015), this regrowth may prevent sampling of the oldest post-collapse 1014 1015 581 products and inhibit estimates of eruption rates and volumes. For example, the relatively 1016 582 small 1956 collapse scar at Bezymianny has been filled by dome lavas within a few decades 1017 1018 583 (Girina, 2013), representing a rapid extrusion rate that would be unresolvable for more 1019 584 ancient events. 1020 1021 585 1022 1023 586 Dating uncertainties also hinder discrimination between rapid (decadal to centennial) and 1024 587 much more gradual shifts in behaviour. There are numerous examples of major collapses 1025 1026 588 where data are too limited to unambiguously correlate specific episodes of volcanism with 1027 589 the timing of collapse (e.g., Galunggung, Indonesia (Bronto, 1989)), or where available data 1028 1029 590 span a restrictive stratigraphic range (e.g., Avachinsky, Kamchatka (Ponomareva et al., 1030 1031 591 2006)) or are coarsely resolved (e.g., Orizaba, Mexico (Schaaf and Carrasco-Núñez, 2010); 1032 592 St Lucia, Lesser Antilles (Boudon et al., 2013); Baru, Panama (Sherrod et al., 2007)). For the 1033 1034 593 vast majority of identified sector collapses in Fig. 1, stratigraphic resolution and age 1035 594 reconstructions do not enable a robust evaluation of how collapse affected subsequent 1036 1037 595 volcanism. The examples discussed here (summarised in Table 3) are thus limited to a small 1038 596 number of better studied volcanoes. These are also geographically skewed, with detailed 1039 1040 597 studies available from many Andean volcanoes and a notable absence of data from Japan and 1041 1042 598 Indonesia. 1043 599 1044 1045 600 7.1. Anomalous post-collapse eruptions 1046 601 If eruptible magma exists beneath a volcano, regardless of whether an incipient eruption 1047 1048 602 caused the collapse, collapse-driven depressurisation may result in magma ascent or other 1049 1050 603 perturbation of the magma reservoir (cf. Table 1). This is supported by the post-collapse 1051 604 eruption of unusually voluminous or compositionally anomalous lavas at several volcanoes. 1052 1053 605 For example, the 12 km 3 collapse of Chimborazo, Ecuador at 60–65 ka was followed by the 1054 606 eruption of homogeneous andesitic lava flows that extend to distances of 22 km and have a 1055 1056 607 total volume of 1–1.5 km 3 (Samaniego et al., 2012). The Holocene collapse of Antuco, Chile, 1057 1058 608 presents a comparable example, where two unusually thick and far-reaching lava flows 1059 609 directly overlie the debris avalanche deposit (Fig. 5). The Sr-isotope composition of these 1060 1061 18 1062 1063 1064 1065 1066 610 lavas is more closely aligned with younger post-collapse rocks (unpublished data), but they 1067 611 are less mafic than the basalts that have dominated all subsequent activity at Antuco (Table 3; 1068 1069 612 Martínez et al., 2018). In both cases, the lava-flow volumes are anomalously large in a long- 1070 613 term context of activity at the volcano, and the lava composition is potentially consistent with 1071 1072 614 mixing between a pre-existing magma body and fresh, mafic input, or the disruption of a 1073 1074 615 previously stable (e.g. vertically zoned or compartmentalised) reservoir. 1075 616 1076 1077 617 Other examples of anomalously voluminous post-collapse effusive eruptions may include 1078 618 Fuya Fuya (Robin et al., 2009) and Tungurahua I (Hall et al., 1999), both in Ecuador. The 1079 1080 619 latter edifice experienced a large but poorly-exposed collapse at the end of its history, 1081 1082 620 followed by emplacement of voluminous dacite lavas, unusually silicic in the long-term 1083 621 history of the volcano. The andesites erupted during the regrowth of the younger Tungurahua 1084 1085 622 II edifice are geochemically distinct from those of Tungurahua I, indicating establishment of 1086 623 a discrete plumbing system (Hall et al., 1999). Tungurahua II itself collapsed at ~3 ka, and 1087 1088 624 the volcano provides a rare example where collapse-associated shifts in behaviour are 1089 1090 625 replicated over more than one eruptive cycle. The Tungurahua II collapse was followed by 1091 626 voluminous (6-km long) and petrologically distinctive (olivine, two-pyroxene and amphibole- 1092 1093 627 phyric) dacite lava flows, and then a pause in volcanism of ~700 years before rapid regrowth 1094 628 of the andesitic Tungurahua III edifice (Hall et al., 1999). A comparable pattern occurs at San 1095 1096 629 Pedro, Chile, where a mid-Holocene collapse was followed by an unusual composite lava 1097 630 flow with a total volume (0.8 km 3) several times larger than younger summit lavas (Costa and 1098 1099 631 Singer, 2002). The composite flow, representing the earliest post-collapse rocks, differs from 1100 1101 632 both older and younger lavas in containing hornblende phenocrysts. As well as having a bulk 1102 633 composition more felsic than any other Holocene lavas at San Pedro, it contains gabbroic 1103 1104 634 xenoliths and basaltic inclusions that extend to more mafic compositions than other Holocene 1105 635 products (Fig. 5). Costa and Singer (2002) interpret these lavas as representing rapid 1106 1107 636 withdrawal from a zoned upper-crustal reservoir, and infer that their textural complexity 1108 1109 637 developed prior to collapse. However, an alternative explanation is that mixing was induced 1110 638 by unloading, and an absence of intervening material or erosion suggests that the 1111 1112 639 emplacement of these complex lavas occurred shortly after sector collapse. 1113 640 1114 1115 641 Post-collapse pyroclastic deposits at Nevado de Colima, Mexico may provide an additional 1116 1117 642 example of decompression-driven disruption of a zoned magma reservoir (Fig. 5). This ~7 1118 643 km 3 collapse is directly overlain by deposits containing juvenile clasts of mixed composition, 1119 1120 19 1121 1122 1123 1124 1125 644 including the most mafic material known from the volcano (although more mafic rocks occur 1126 645 at outlying scoria cones) (Robin et al., 1990). At both Colima and San Pedro, compositions 1127 1128 646 that are unknown at the surface in other intervals are observed in the earliest post-collapse 1129 647 rocks, but not in subsequent eruptions. The hornblende-phyric pumice erupted after the 1888 1130 1131 648 collapse of Ritter Island, which appears to have otherwise erupted relatively homogeneous 1132 1133 649 basalts and basaltic andesites (Watt et al., 2019), provides a comparable example, as do the 1134 650 Secche di Lazzaro deposits at Stromboli (Petrone et al., 2009; see Section 7.2 and Table 3). 1135 1136 651 Collectively, these examples are consistent with anomalous mixing and ascent dynamics 1137 652 following collapse, and highlight that compositions erupted under equilibrium (i.e. without 1138 1139 653 collapse-driven reservoir perturbation) conditions may present an incomplete or biased 1140 1141 654 picture of the compositional range of crustal magmas. 1142 655 1143 1144 656 7.2. Post-collapse regrowth through mafic volcanism 1145 657 A post-collapse shift to more mafic erupted compositions occurs at several volcanoes, 1146 1147 658 maintained over multiple eruptions (Fig. 6). This pattern includes examples with anomalous 1148 1149 659 post-collapse eruptions, such as Antuco, where the voluminous lava flows were followed by 1150 660 more mafic monotonous lavas and scoria deposits that have infilled the collapse scar 1151 1152 661 (representing a higher eruptive flux than pre-collapse activity) and persist to the present day 1153 662 (Martínez et al., 2018). Similarly, at Martinique, the Le Prêcheur collapse is bracketed by two 1154 1155 663 near identical lavas (suggesting post-collapse eruption from an existing magma body), but is 1156 664 followed by a sharp shift to mafic compositions before an eventual return to more evolved 1157 1158 665 magmas (Germa et al., 2011). 1159 1160 666 1161 667 Eruptions of relatively mafic magmas may be maintained for several thousand years 1162 1163 668 following collapse, and this is consistent with conditions that favour rapid crustal transit and 1164 669 the ascent of denser magmas to the surface. At Planchón, Chile, pre-collapse basaltic 1165 1166 670 andesites are followed by post-collapse basalts that lack evidence for upper crustal storage 1167 1168 671 (Tormey et al., 1995; Tormey, 2010) and are compositionally similar to the oldest pre- 1169 672 collapse rocks, thus supporting a general cyclical coupling of edifice growth and magma 1170 1171 673 evolution, with increased loading favouring upper crustal stalling of denser, mafic magmas 1172 674 (e.g., Schindlbeck et al., 2014). Alongside more mafic compositions and evidence of reduced 1173 1174 675 storage and assimilation, an elevated eruptive flux is observed in several examples. At 1175 1176 676 Tungurahua, rapid regrowth of the TIII edifice (following the dacite lava eruptions and a 700- 1177 677 year pause in activity) was achieved by volumetric growth rates of ~1.5 km 3/kyr that have 1178 1179 20 1180 1181 1182 1183 1184 678 persisted since 2.3 ka (Hall et al., 1999). The early stages of this cone rebuilding produced 1185 679 homogeneous basaltic andesites, suggesting relatively rapid crustal transit, but recent magmas 1186 1187 680 are more evolved and may hint at a transition towards more prolonged shallow storage. A 1188 681 broadly similar pattern is suggested by the evolution of Ollagüe, Chile, following a relatively 1189 1190 682 smaller flank collapse (Feeley et al., 1993; Clavero et al., 2004), as well as the examples 1191 1192 683 shown in Fig. 6. 1193 684 1194 1195 685 At Parinacota, Chile, the 6 km 3 collapse at 8.8 ka (Jicha et al., 2015) was followed by 1196 686 extremely high eruptive flux of up to 10 km 3/kyr for one to two thousand years, sufficient to 1197 1198 687 entirely bury the collapse scar. Post-collapse rocks are relatively homogeneous and more 1199 1200 688 mafic than pre-collapse lavas (Wörner et al., 1988; Hora et al., 2007, 2009), including the 1201 689 most primitive compositions known at the volcano (Fig. 6). The collapse appears to have 1202 1203 690 occurred within a long-term trend towards more mafic compositions, but regardless of the 1204 691 compositional impact of collapse, the increased post-collapse flux of relatively monotonous 1205 1206 692 magmas represents a sharp change from pre-collapse activity. Evidence for long-term upper- 1207 1208 693 crustal stagnation under pre-collapse conditions contrasts with much more limited evidence 1209 694 of assimilation in younger products (Hora et al., 2009). A spatial change in the post-collapse 1210 1211 695 plumbing system, with a diminished upper crustal reservoir, is also suggested by plagioclase 1212 696 zoning patterns (Ginibre and Wörner, 2007). 1213 1214 697 1215 698 Stromboli, Italy, provides a further example of a coupled post-collapse shift in magma 1216 1217 699 chemistry and eruptive behaviour (Francalanci et al., 1989; Bertagnini and Landi, 1996; 1218 1219 700 Tibaldi, 2001, 2004; Petrone et al., 2009; Vezzoli et al., 2014). The 14 ka Upper Vancori 1220 701 collapse (2.2 ± 0.9 km 3) ended a period erupting variable basic and intermediate magmas 1221 1222 702 (Fig. 6), and was followed by more compositionally restricted mafic volcanism (Francalanci 1223 703 et al., 1989; Vezzoli et al., 2014) at a higher volumetric growth rate (Tibaldi, 2004; Vezzoli et 1224 1225 704 al., 2014). These post-collapse magmas are unusually potassic, with relatively monotonous 1226 1227 705 bulk compositions and petrological characteristics. However, their variable Sr-isotope 1228 706 compositions suggests differential assimilation during crustal transit, coincident with the 1229 1230 707 reduction of the upper crustal reservoir (Francalanci et al., 1989; Hornig-Kjarsgaard et al., 1231 708 1993). The unstable post-collapse cone at Stromboli has failed to achieve the dimensions of 1232 1233 709 the preceding edifice, and has experienced three younger and relatively closely-spaced 1234 1235 710 collapses from ~6 ka (each <50% of the volume of the Upper Vancori collapse; Tibaldi, 1236 711 2001, 2004), maintaining high growth rates but with a further compositional shift to basaltic 1237 1238 21 1239 1240 1241 1242 1243 712 shoshonitic magmas following the ~6 ka collapse (Fig. 6), and a predominance of explosive 1244 713 over effusive activity (Francalanci et al., 1989; Hornig-Kjarsgaard et al., 1993)). The 1245 1246 714 Neostromboli eruptive phase, which followed the Upper Vancori collapse and was terminated 1247 715 by a ~1km3 sector collapse at 6 ka (Tibaldi, 2001), provides additional support to a cyclic 1248 1249 716 model of post-collapse regrowth through mafic volcanism and the re-establishment of a more 1250 1251 717 evolved, upper crustal reservoir (Petrone et al., 2009; Vezzoli et al., 2014). The youngest 1252 718 magmas erupted in this 8 kyr period are also the most evolved Neostromboli rocks (Fig. 6), 1253 1254 719 and derived from fractional crystallisation of the more primitive magmas that fed earlier post- 1255 720 collapse eruptions (Vezzoli et al., 2014). These biotite-shoshonites also fed powerful 1256 1257 721 phreatomagmatic explosive eruptions associated with the 6 ka sector collapse (Petrone et al., 1258 1259 722 2009), whereas all subsequent volcanism has been more primitive. Thus, since the 1260 723 Neostromboli sector collapse at 6 ka, the upper-crustal reservoir does not appear to have been 1261 1262 724 re-established (Vezzoli et al., 2014). The phreatomagmatic activity that accompanied the 1263 725 Neostromboli collapse (the Secche di Lazzaro eruptions) erupted petrographically 1264 1265 726 heterogeneous products, suggestive of disruption of a complex, stratified shallow plumbing 1266 1267 727 system, and thus consistent with the processes described in Section 7.1 (Petrone et al., 2009). 1268 728 Based on mineral zoning, compositional variability and disequilibrium textures, Petrone et al. 1269 1270 729 (2009) rule out fresh mafic input as a trigger for the Secche di Lazzaro eruptions (and by 1271 730 implication, as a trigger for the Neostromboli collapse); rather, physical and petrological 1272 1273 731 observations are consistent with decompression-driven eruption of shallowly-stored magma 1274 732 as a direct result of the collapse (Bertagnini and Landi, 1996; Petrone et al., 2009). The 1275 1276 733 evolution of Stromboli between the Upper Vancori and Neostromboli collapses thus 1277 1278 734 encapsulates the range of processes described in Sections 7.1 to 7.3, across an 8 kyr period 1279 735 (Fig. 6). 1280 1281 736 1282 737 7.3. Re-establishment of upper crustal storage 1283 1284 738 Longer timescale reconstructions suggest that the mafic activity at volcanoes such as 1285 1286 739 Stromboli, Parinacota and Antuco may represent a temporary cone-rebuilding stage. Edifice 1287 740 regrowth and increased loading may eventually promote upper-crustal storage, manifested as 1288 1289 741 a return to more evolved erupted compositions. The ~32 ka collapse of Pelée, Martinique, 1290 742 initiated a similar pattern to that of the older Le Prêcheur collapse (Germa et al., 2011): pre- 1291 1292 743 collapse andesites were followed by denser basaltic andesites at an elevated eruption rate 1293 1294 744 (inferred from tephra-deposit frequencies and volumetric reconstructions) (Germa et al., 1295 745 2015). However, within ~10 kyr eruptions returned to predominantly andesitic compositions 1296 1297 22 1298 1299 1300 1301 1302 746 (Boudon et al., 2013) (Fig. 6), but enclaves in these young andesites show that deeper mafic 1303 747 magmas were still feeding the newly established upper crustal reservoir. 1304 1305 748 1306 749 The evolution of Pelée has a counterpart in the basaltic South Soufrière Hills episode on 1307 1308 750 Montserrat (Cassidy et al., 2015a), which followed the 130 ka D2 collapse and is 1309 1310 751 compositionally anomalous in the long-term history of this andesitic island (Fig. 6). Although 1311 752 the duration of South Soufrière Hills volcanism is difficult to constrain, Cassidy et al. (2015a) 1312 1313 753 suggest that the onset of post-collapse mafic volcanism was rapid (<100 years). Available 1314 754 dates for South Soufrière Hills are also identical within error (Harford et al., 2002), 1315 1316 755 suggesting elevated eruption rates with a maximum duration of a few thousand years. Cassidy 1317 1318 756 et al. (2015a) interpret the basalts to have risen rapidly from mid-crustal depths, suggesting 1319 757 an absent or inactive upper-crustal storage system in the early post-collapse period, in 1320 1321 758 common with the other examples cited above. 1322 759 1323 1324 760 Activity at Montserrat returned to andesitic dome-forming eruptions after the South Soufrière 1325 1326 761 Hills episode, erupting rocks similar to pre-collapse lavas but distinctive in containing 1327 762 phenocryst hornblende (this distinction has persisted to the present day) (Harford et al., 2002; 1328 1329 763 Cassidy et al., 2015a). This pattern is replicated at some Andean volcanoes (de Silva et al., 1330 764 1993), but is the reverse to that observed at Parinacota, where the dominant mineralogy 1331 1332 765 changes from a pre-collapse hornblende andesite to a two-pyroxene assemblage (Wörner et 1333 766 al., 1988). The direction of the change is not necessarily significant, but simply highlights 1334 1335 767 that the re-established crustal plumbing system was distinct in terms of its predominant 1336 1337 768 pressure and temperature conditions from that which existed previously. 1338 769 1339 1340 770 7.4. Long-term collapse-induced shifts in storage and plumbing systems 1341 771 The impacts of large-scale edifice collapse are not always manifested through more mafic 1342 1343 772 erupted compositions of the type described above. This may be because the post-collapse 1344 1345 773 stress regime, although modified from preceding conditions, still promotes upper crustal 1346 774 stalling of denser magmas. In such cases, evidence of enhanced ascent of mafic, lower-crustal 1347 1348 775 magmas may be more subtle. At San Pedro, post-collapse replenishment (following the 1349 776 anomalous lava flows; Fig. 5) led to eruptions of basaltic andesites with similar major- 1350 1351 777 element chemistry to pre-collapse lavas, but with distinct trace-element and isotope 1352 1353 778 signatures (Costa and Singer, 2002). This implies a crustal processing history that is 1354 779 temporally distinct from the pre-collapse magmas. The same conclusion can be drawn at 1355 1356 23 1357 1358 1359 1360 1361 780 Chimborazo (Fig. 7; Table 3), Tungurahua (Hall et al., 1999) and elsewhere, where 1362 781 permanent changes in bulk compositional trends support the establishment of a discrete 1363 1364 782 plumbing system, overwriting any influence of older magmatic contributions. There are many 1365 783 other less well studied volcanoes where available data nevertheless hint at post-collapse 1366 1367 784 modification of the plumbing system. Examples include the presence of unusually mafic 1368 1369 785 inclusions in post-collapse lava domes at Tata Sabaya, Bolivia (de Silva et al., 1993), and a 1370 786 change to predominantly effusive activity after the sector collapse of Santa Ana, El Salvador 1371 1372 787 (Siebert et al., 2004). 1373 788 1374 1375 789 The very large late-Pleistocene collapse of Shiveluch marks a transition in both chemistry 1376 1377 790 and eruption style (Table 3). Although pre- and post-collapse rocks span a similar range of 1378 791 silica contents, the post-collapse mafic magmas are more primitive (e.g., in terms of Mg#; 1379 1380 792 Table 3) than older equivalents, thus supporting renewed mafic input (Gorbach et al., 2013) 1381 793 (Fig. 7). Compositional trends suggest that pre-collapse magmas predominantly evolved via 1382 1383 794 fractional crystallisation, while post collapse magmas are consistent with mixing between 1384 1385 795 deeper basalts and shallower felsic magmas, interpreted by Gorbach et al. (2013) as 1386 796 indicating a simpler and smaller shallow reservoir in the post-collapse period. Belousov et al. 1387 1388 797 (1999) suggest that these changes had additional consequences for eruption style and edifice 1389 798 stability, producing intermediate magmas with higher viscosity in the post-collapse period, 1390 1391 799 emplaced as lava domes. These built a steep edifice prone to destabilisation upon subsequent 1392 800 intrusion, and susceptible to repeated smaller-scale collapses of the type observed in 1964. In 1393 1394 801 this respect, post-collapse behaviour at Shiveluch mirrors that at Stromboli, over a 1395 1396 802 comparable time period (cf. Tibaldi, 2004). 1397 803 1398 1399 804 Observations at several volcanoes suggest that sector collapse focuses shallow dyke 1400 805 propagation towards the collapse scar (Tibaldi et al., 2008), with the vent migrating to the 1401 1402 806 centre of the collapse amphitheatre at volcanoes such as Planchón (Tormey, 2010), Ollagüe 1403 1404 807 and Stromboli (Tibaldi et al., 2008) (comparable processes have also been suggested at 1405 808 ocean-island volcanoes; Maccaferri et al., 2017). Post-collapse rebuilding often buries the 1406 1407 809 scar (e.g., Parinacota), and is generally centred on a similar position to the pre-collapse 1408 810 conduit. Thus, assuming broadly vertical storage geometries, younger magmas appear to pass 1409 1410 811 through preceding storage regions while retaining a discrete chemical signature. The South 1411 1412 812 Soufrière Hills episode at Montserrat is unusual in lying to one side of the collapse scar (cf. 1413 813 Cassidy et al., 2015b), but given observations elsewhere there is no general requirement for 1414 1415 24 1416 1417 1418 1419 1420 814 post-collapse mafic magmas to bypass upper crustal reservoirs to reach the surface. Indeed, 1421 815 by reaching the surface, the eruption of deeper, mafic magmas implies the absence of a 1422 1423 816 mobile upper crustal reservoir under early post-collapse conditions. 1424 817 1425 1426 818 7.5. A coherent model? 1427 1428 819 The sections above provide several examples where sector collapses mark clear transitions in 1429 820 erupted magma compositions or style of volcanic activity, but these effects are not identical 1430 1431 821 at each volcano. Nevertheless, the observations can fit a single coherent model if inter- 1432 822 volcano differences are attributable to local factors. For example, the volume of a sector 1433 1434 823 collapse, in conjunction with storage depths, reservoir volumes, and the presence of eruptible 1435 1436 824 magma, is likely to determine the observed effect. Notwithstanding these differences, a 1437 825 consistent model emerges from the examples above, summarised in Fig. 8. 1438 1439 826 1440 827 7.5.1. Collapse-driven reservoir disruption, or vice versa? 1441 1442 828 Determining the causal sequence of collapse-associated transitions in volcanic behaviour is 1443 1444 829 challenging. If it is fresh magma input that drives edifice instability, then shifts in magma 1445 830 compositions may simply be temporally associated with sector collapse, but not initiated by 1446 1447 831 the unloading process. However, although it is not possible to definitively exclude this 1448 832 interpretation, large-scale sector collapse provides a plausible mechanism for disruption and 1449 1450 833 reconfiguration of an otherwise stable plumbing system (and one that has developed in 1451 834 equilibrium with the overlying load), whereas no specific mechanism exists for the reverse. 1452 1453 835 The anomalous eruptions described in Section 7.1 followed but did not accompany collapse 1454 1455 836 (i.e. they are not present as syn-collapse pyroclastic deposits or juvenile debris-avalanche 1456 837 blocks), suggesting that magma ascent did not precede or trigger collapse in these instances. 1457 1458 838 Petrone et al. (2009) present a rare example where detailed mineralogical analyses are used to 1459 839 specifically exclude the possibility of fresh magmatic input prior to the Neostromboli 1460 1461 840 collapse, implying that the intense phreatomagmatic eruptions associated with the collapse 1462 1463 841 directly reflect decompression-driven disruption of the shallow system, and supporting the 1464 842 above interpretation. The replication of compositional anomalies in several examples, 1465 1466 843 contemporaneous with collapse but absent in both older and younger rocks, lends further 1467 844 support for a discrete external process (i.e. sector collapse) driving this atypical behaviour, 1468 1469 845 and one that falls outside the prevailing factors that govern magma storage and eruption at 1470 1471 846 other times in a volcano’s history. This argument is also supported by the sharpness of the 1472 847 compositional shifts between pre- and post-collapse magmas (Figs. 6 and 7), which imply the 1473 1474 25 1475 1476 1477 1478 1479 848 development of a structurally distinct plumbing system during post-collapse regrowth. The 1480 849 interpretation presented here suggests that erupted magma compositions spanning sector 1481 1482 850 collapses are governed by upper crustal processes that reflect the interaction between edifice 1483 851 loads and the stability and growth of shallow magma reservoirs. They do not require (but nor 1484 1485 852 do they exclude) any variation in the flux or ascent of deeper, primitive magmas, but the 1486 1487 853 capacity of such magmas to reach the surface, rather than being captured within shallow 1488 854 reservoirs, may vary as a result of the prevailing upper crustal conditions. 1489 1490 855 7.5.2. Persistent impacts on plumbing systems and cyclic development 1491 856 Although the precise age and duration of the compositionally anomalous eruptions described 1492 1493 857 above are not generally constrained, their stratigraphic position directly overlying debris 1494 1495 858 avalanche deposits or pre-collapse rocks (e.g., Ritter, Stromboli, Antuco; Table 3), generally 1496 859 as a single eruptive unit, suggests that they represent short-lived events that potentially 1497 1498 860 occurred immediately after collapse. In contrast, subsequent periods of more persistent mafic 1499 861 volcanism may span time intervals of several thousand years (Fig. 6). This suggests that the 1500 1501 862 compositionally anomalous eruptions mark the death of a plumbing system that is out of 1502 1503 863 equilibrium with the collapse-modified surface load. That anomalous eruptions are not 1504 864 observed after all large sector collapses may simply reflect that eruptible magma wasn’t 1505 1506 865 present at the time of collapse (Fig. 8), or that post-collapse stress conditions were not 1507 866 favourable to magma ascent (Table 1) – one of the outcomes predicted by Pinel and Albino 1508 1509 867 (2013). The former interpretation is more consistent with observations: if the upper crustal 1510 868 reservoir is retained after collapse, but magma ascent is simply temporarily hindered, then 1511 1512 869 this can’t explain long-term changes in erupted compositions. The fact that sharp 1513 1514 870 compositional changes are observed in multiple examples, coeval with collapse, implies more 1515 871 fundamental impacts on plumbing systems, leading to a post-collapse magmatic regime that 1516 1517 872 is discrete from the preceding one. 1518 873 1519 1520 874 The eruption sequences at San Pedro, Tungurahua, Antuco, Stromboli and elsewhere can thus 1521 1522 875 be interpreted in a framework of a pre-existing mobile reservoir, where unloading drives 1523 876 mixing of a previously stagnant magma body. In this interpretation, the earliest post-collapse 1524 1525 877 magmas would not have erupted without sector collapse. It is notable that several of these 1526 878 anomalous eruptions were effusive (Hall et al., 1999; Costa and Singer, 2002; Martínez et al., 1527 1528 879 2018), and not necessarily initiated by high overpressures or the ascent of volatile-rich 1529 1530 880 magmas from the lower crust, but simply by upper crustal decompression (cf. Petrone et al., 1531 881 2009). 1532 1533 26 1534 1535 1536 1537 1538 882 1539 883 Other examples, such as Soufrière Hills, Pelée and Tungurahua, support a coupling between 1540 1541 884 edifice growth and magma storage, whereby post-collapse eruptions of deeper, mafic 1542 885 compositions are followed by a return to more evolved compositions (Hall et al., 1999; 1543 1544 886 Boudon et al., 2013; Cassidy et al., 2015a) upon partial edifice regrowth and the re- 1545 1546 887 establishment of upper-crustal stalling of denser, mafic magmas. Longer timescale 1547 888 observations thus support a broad cyclicity of mafic to evolved volcanism, delineated by 1548 1549 889 collapses, which are followed by sharp shifts to more mafic erupted compositions. This 1550 890 provides strong evidence that collapse is followed by the establishment of a storage and 1551 1552 891 plumbing system distinct from the preceding one and, by implication, supports an inherent 1553 1554 892 coupling between surface loading and the factors controlling magma storage and ascent. 1555 893 1556 1557 894 The observation of sharp and wholesale shifts in magma composition calls into question 1558 895 models of magma storage that propose long-lived and large-volume liquid bodies, because it 1559 1560 896 is difficult to conceive how such bodies could be bypassed or sufficiently diluted by younger 1561 1562 897 magmas to erase preceding compositional signatures. Rather, reservoirs dominated by crystal 1563 898 mushes, susceptible to becoming immobile over short timescales, are more consistent with 1564 1565 899 observations that require rapid adjustment of upper crustal plumbing systems to enable 1566 900 changes in dominant storage depths and the transit of magmas with persistently different 1567 1568 901 assimilation and crystallisation histories. 1569 902 1570 1571 903 7.5.3. Counter-examples and historical collapses 1572 1573 904 Many arc volcanoes affected by sector collapse have only limited age and compositional data, 1574 905 and even examples with better stratigraphic constraints, such as those in Fig. 6, may suffer 1575 1576 906 from data gaps of many thousands of years or unresolvable age relationships. An absence of 1577 907 mafic post-collapse magmas (e.g., de Silva et al., 1993), or even an apparent shift to more 1578 1579 908 felsic post-collapse compositions (e.g. Boudon et al., 2013), may thus be attributable to 1580 1581 909 sparse stratigraphic reconstructions (cf. collapses on Martinique and St Lucia, much older 1582 910 than those outlined above, where the time gap between collapse and the next dated eruption is 1583 1584 911 poorly resolved; Boudon et al., 2013), and may also be affected by burial of the earliest post- 1585 912 collapse magmas. 1586 1587 913 1588 1589 914 At all arc volcanoes with high-resolution stratigraphic reconstructions spanning large-volume 1590 915 (several cubic kilometres; Table 3) sector collapses, the observations are consistent with the 1591 1592 27 1593 1594 1595 1596 1597 916 processes outlined in Fig. 8. However, with the exception of Ritter Island in 1888 (Watt et al., 1598 917 2019), historical sector collapses show no such dramatic impacts on the underlying magma 1599 1600 918 system. This may simply reflect the small magnitude of historical events (Table 2) relative to 1601 919 those in Table 3. The 1964 collapse at Shiveluch, for example, involved a fraction of the 1602 1603 920 material mobilised in its late-Pleistocene collapse. If collapse-induced stress changes are 1604 1605 921 comparable to those experienced by crustal magma reservoirs during normal cycles of 1606 922 pressurisation and eruption, then the plumbing system may withstand the impacts of collapse 1607 1608 923 without major modification. Similar behaviour is observed in general eruptive activity: small 1609 924 eruptions fit within coherent trends of composition and style, while major eruptions, 1610 1611 925 involving several cubic kilometres of magma, may mark sharp transitions in behaviour (e.g. 1612 1613 926 Schindlbeck et al., 2014; Gavrilenko et al., 2016). What is notable about sector collapse is 1614 927 that the magma plumbing system may be modified without evacuation of the upper crustal 1615 1616 928 reservoir. This implies that material within the pre-collapse reservoir remains in the crust, but 1617 929 a new plumbing system, giving rise to distinct compositions and evolutionary trends, can 1618 1619 930 overprint the former regime. 1620 1621 931 1622 932 Taranaki, New Zealand, provides an example of a volcano that may not fit into the scheme 1623 1624 933 outlined above. Although Taranaki’s rocks display a gradual evolution to more evolved and 1625 934 more potassic compositions over its 130 kyr lifetime (Zernack et al., 2012), available data 1626 1627 935 show no apparent sharp changes in magma composition following collapse (post-collapse 1628 936 stratigraphic constraints are, however, absent beyond the Holocene, due to burial during 1629 1630 937 younger regrowth cycles). The volcano is remarkable for the number of identified collapses, 1631 1632 938 and although their precise volume is difficult to constrain (Zernack et al., 2009, 2011), 1633 939 several appear to be comparable in scale to the events in Table 3 (Zernack et al., 2012). The 1634 1635 940 collapse frequency is not as high as that at Augustine (Béget and Kienle, 1992) and Shiveluch 1636 941 (Belousov et al., 1999; Ponomareva et al., 2006), but the larger average collapse volume 1637 1638 942 suggests that Taranaki, like these volcanoes, is characterised by a notably high magma flux 1639 1640 943 and edifice growth rate. This elevated flux may limit the coupling between surface loads and 1641 944 the upper crustal plumbing system, if edifice growth outpaces the crustal response. This is 1642 1643 945 suggested by the relatively constant long-term effusive flux at Taranaki (Zernack et al., 1644 946 2012), which contrasts with the more pulsatory eruptive flux observed at many volcanoes. 1645 1646 947 (Fig. 4). 1647 1648 948 1649 949 8. Collapses at intraplate ocean islands 1650 1651 28 1652 1653 1654 1655 1656 950 Intraplate ocean islands represent a second volcano-tectonic setting where several 1657 951 comprehensive studies have identified large-scale flank collapses (e.g., Moore et al., 1989; 1658 1659 952 Masson et al., 2002, 2008), and where the question of how collapse affects magma generation 1660 953 and storage has been addressed by different authors (cf. Longpré et al., 2009; Manconi et al., 1661 1662 954 2009; Boulesteix et al., 2012; Hunt et al., 2018). Despite the common physical process of 1663 1664 955 gravitational unloading, the magmatic impacts of ocean-island collapses should be evaluated 1665 956 separately to subduction zone volcanoes for a number of reasons. First, ocean islands are 1666 1667 957 dominated by mafic volcanism and do not typically develop long-lived upper-crustal 1668 958 plumbing systems characterised by intermediate to evolved magmas (even if shallow mafic 1669 1670 959 reservoirs develop; cf. Clague, 1987; Frey et al., 1991; Amelung and Day, 2002; Galipp et al., 1671 1672 960 2006; Hildner et al., 2012). With limited shallow magma storage and without evolved 1673 961 compositions, the influence of loading on magma plumbing systems may be more poorly 1674 1675 962 developed than in arc settings. This is not true of all ocean islands, however, with locations 1676 963 such as Tenerife, Canary Islands, producing large explosive eruptions of felsic magma (e.g., 1677 1678 964 Bryan et al., 1998). Second, the dimensions of ocean islands (Figure 1) are in some cases 1679 1680 965 comparable to the elastic thickness of the underlying (oceanic) lithosphere, and the 1681 966 dimensions of collapse-driven mass redistribution may even be significant in this context. 1682 1683 967 Unlike arc settings, it cannot be assumed that the island mass is supported by the strength of 1684 968 the lithosphere, and there may be an isostatic response to landslide mass redistribution (cf. 1685 1686 969 Smith and Wessel, 2000), and potentially an influence on melt production through mantle 1687 970 decompression (e.g., Presley et al., 1997). Finally, there are variable spatial relationships 1688 1689 971 between flank collapses on ocean islands and the underlying plumbing system. The aspect 1690 1691 972 ratio of collapses is generally lower (Figure 1), despite very large collapse volumes, and 1692 973 collapses may not necessarily overlie the central plumbing system due to the rift-zone 1693 1694 974 structure that characterises many ocean islands (cf. the Orotava and Güímar collapses on 1695 975 Tenerife (Fig. 9), which lie on a rift zone outside the central Cañadas caldera structure; Martí 1696 1697 976 et al., 1997; Carracedo et al., 2011). 1698 1699 977 1700 978 8.1. Evidence of plumbing system modifications 1701 1702 979 Pressure constraints at several ocean islands suggest pre-eruptive magma storage at upper 1703 980 mantle depths, but observations imply that such deep plumbing systems can nevertheless be 1704 1705 981 directly influenced by sector collapse (Table 4). At Fogo, Cape Verde Islands, magmas 1706 1707 982 erupted after the Monte Amarelo collapse (123-62 ka) equilibrated at greater depths (~25 km) 1708 983 than pre-collapse magmas (~18 km), which showed a long-term shallowing trend prior to 1709 1710 29 1711 1712 1713 1714 1715 984 collapse (although the absolute timescale of this is not well constrained; Hildner et al., 2012) 1716 985 (Fig. 10). Subsequent post-collapse regrowth involved magmas with a shallower and broader 1717 1718 986 range of storage depths, suggesting development of an increasingly complex crustal plumbing 1719 987 system. A broadly similar pattern, spanning comparable depth ranges, has been identified at 1720 1721 988 La Palma, Canary Islands, where the Bejenado volcano grew rapidly after the Cumbre Nueva 1722 1723 989 collapse (560 ka; Guillou et al., 1998; Galipp et al., 2006) (Fig. 10). Younger Bejenado 1724 990 magmas show evidence of increased fractional crystallisation and possibly decreasing magma 1725 1726 991 supply rates as the edifice developed, and the range of storage depths suggests entirely 1727 992 distinct pre- and post-collapse plumbing systems (Galipp et al., 2006). 1728 1729 993 1730 1731 994 A detailed evaluation of shield-basalt sequences at the Teno massif, Tenerife, highlights 1732 995 further distinctive shifts in volcanism (Longpré et al., 2009) associated with two >20-25 km 3 1733 1734 996 landslides. These landslides interrupt and sharply reverse long-term trends towards more 1735 997 silicic and less magnesian lavas, and appear to have disrupted the pre-existing plumbing 1736 1737 998 system. They resulted in explosive eruptions from minor shallow reservoirs (unusual in the 1738 1739 999 long-term effusion-dominated context of the shield) and enabled the ascent and eruption of 1740 1000 deep, dense, crystal-rich mafic magmas in the early post-collapse period. Similar patterns 1741 1742 1001 follow the El Golfo collapse on El Hierro, Canary Islands (Manconi et al., 2009), where the 1743 1002 pre-collapse stratigraphy includes evolved compositions (trachytes). Such rocks are absent in 1744 1745 1003 the post-collapse stratigraphy, which is dominated by mafic, crystal-rich lavas. Numerical 1746 1004 models (Manconi et al., 2009) suggest that pressure changes in the deep plumbing system, on 1747 1748 1005 the order of 1 MPa, are sufficient to disrupt stored magma, drive volumetric expansion and 1749 1750 1006 potentially initiate fracturing and dyke ascent. 1751 1007 1752 1753 1008 Volcan Ecuador, in the Galapagos, provides a further example where sector collapse is 1754 1009 associated with a shift in the locus of volcanism following disruption of the shallow plumbing 1755 1756 1010 system (comparable to changes in vent distribution following the SW landslides on Mauna 1757 1758 1011 Loa, Hawaii (Lipman et al., 1990)), and where post-collapse magmas have elevated MgO 1759 1012 contents, suggesting favourable ascent of mafic melts (Geist et al., 2002). Shifts in vent 1760 1761 1013 location, associated with modified magma ascent paths under post-collapse conditions, have 1762 1014 been observed in many ocean island (and some arc) settings (cf. Maccaferri et al., 2017) and 1763 1764 1015 may play a role in the development of discrete post-collapse plumbing systems. A direct 1765 1766 1016 relationship between rift-zone configuration and flank stability has been proposed at La 1767 1017 Palma (Day et al., 1999a), and on Tenerife it has been proposed that rift-zone reorganisation 1768 1769 30 1770 1771 1772 1773 1774 1018 is a direct consequence of large-scale flank collapses (Walter et al., 2005), which may also 1775 1019 promote more centralised and evolved subsequent volcanism (Carracedo et al., 2011). 1776 1777 1020 Maccaferri et al. (2017) propose that post-collapse stress redistribution, focusing dyke ascent 1778 1021 towards the collapse scar, can potentially act as a feedback mechanism promoting edifice 1779 1780 1022 regrowth and subsequent collapses in the same area. Such a mechanism may be particularly 1781 1782 1023 influential following very large volume collapses in a neutral tectonic environment, as 1783 1024 characterised by many ocean-island volcanoes. As suggested by the evolution of the NE rift 1784 1785 1025 zone on Tenerife, focusing of volcanism within collapse scars can ultimately promote 1786 1026 shallow magma storage, differentiation, and the eruption of more evolved magma 1787 1788 1027 compositions (Carracedo et al., 2011). Thus, the structural as well as the magmatic evolution 1789 1790 1028 of islands such as Fogo (Day et al., 1999b; Maccaferri et al., 2017) and Tenerife (Carracedo 1791 1029 et al., 2011) may directly reflect the past history of sector collapse. 1792 1793 1030 The above examples indicate that although storage conditions and depth ranges may differ 1794 1031 from arc settings, the cyclic pattern of plumbing system modification at ocean-island 1795 1796 1032 volcanoes, characterised by the transit of deeper magmas via simpler ascent routes in the 1797 1798 1033 post-collapse period, with concomitant shifts in chemistry, petrography, and possibly 1799 1034 eruption style and vent position (Table 4), is comparable to the cone-rebuilding phase in Fig. 1800 1801 1035 8. This implies that collapse-initiated reorganisation of magmatic plumbing systems is a 1802 1036 general process. 1803 1804 1037 1805 1038 8.2. Impacts on shallow magma reservoirs 1806 1807 1039 Tenerife is unusual among ocean islands in having a central, large-volume shallow reservoir, 1808 1809 1040 dominated by evolved compositions (Carracedo et al., 2007) (Fig. 9). Indeed, large-scale 1810 1041 sector collapses have been cited as a possible reason for the absence of shallow reservoirs on 1811 1812 1042 some other ocean islands (Amelung and Day, 2002; Geist et al., 2002). The cyclic Bandas del 1813 1043 Sur ignimbrite deposits on Tenerife (Bryan et al., 1998, 2002; Brown et al., 2003; Edgar et 1814 1815 1044 al., 2007) show that evolved magmas have been erupted for prolonged periods of the island’s 1816 1817 1045 recent history, and the impact of the Icod collapse (~175 ka) on the shallow reservoir 1818 1046 (Carracedo et al., 2007; Boulesteix et al., 2012) provides further evidence that large ocean 1819 1820 1047 island landslides impact magma plumbing systems in an essentially identical way to that 1821 1048 observed in arcs. 1822 1823 1049 1824 1825 1050 The Icod landslide deposit is associated with phonolitic pumiceous deposits, but the base of 1826 1051 the collapse scar is infilled by mafic lavas, which were erupted at elevated rates over a 10 kyr 1827 1828 31 1829 1830 1831 1832 1833 1052 period (Boulesteix et al., 2012). Regrowth is interpreted to have led to stalling of mafic 1834 1053 magmas, accompanied by eruption of increasingly evolved compositions at decreasing rates. 1835 1836 1054 This ultimately led to the growth of the phonolitic Teide volcano, fed from an established 1837 1055 upper crustal reservoir (Fig. 9). This sequence replicates very closely the model put forward 1838 1839 1056 in arc settings, and shows the same transitions in magma composition and storage, on a 1840 1841 1057 broadly similar timescale, observed at Montserrat, Martinique and the other examples cited 1842 1058 above (Fig. 8). 1843 1844 1059 1845 1060 It is clear that an evolved upper-crustal reservoir existed at the time of the Icod collapse. 1846 1847 1061 Phonolitic pumice deposits consistent in age and composition with the Abrigo ignimbrite 1848 1849 1062 (Fig. 9) are both cut by the landslide and appear in collapse-associated breccias (Boulesteix et 1850 1063 al., 2012), as well as appearing in the upper units of the Icod turbidite (Hunt et al., 2018) and 1851 1852 1064 tsunami deposits (Paris et al., 2017). This suggests that a large explosive eruption occurred in 1853 1065 the latter stages of collapse, but that similar magma had also fed previous eruptions. Whether 1854 1855 1066 an incipient explosive eruption (i.e. shallow magma ascent) led to collapse, or if the 1856 1857 1067 accompanying eruption was truly collapse-triggered, cannot be deduced. In any case, the next 1858 1068 erupted lavas were mafic and no phonolitic deposits appear in the younger stratigraphy for 1859 1860 1069 tens of thousands of years, replicating the patterns of mafic renewal observed at arc 1861 1070 volcanoes following termination of the upper crustal storage system. 1862 1863 1071 1864 1072 Hunt et al. (2018) suggest that the pattern of contemporaneous pumice clasts occurring in the 1865 1866 1073 upper units of collapse-derived turbidites is replicated by the Orotava collapse (~535 ka) on 1867 1868 1074 Tenerife (possibly linked to the Granadilla eruption at 560-600 ka). A collapse-triggered 1869 1075 eruption seems less likely in this case, given the location of the Orotava landslide outside the 1870 1871 1076 central caldera structure on Tenerife (Fig. 9), but it is plausible that pre-eruptive unrest led to 1872 1077 the collapse. 1873 1874 1078 1875 1876 1079 8.3. A direct influence on melting? 1877 1080 Chemical evidence in support of a direct influence of collapse on mantle melt fraction is 1878 1879 1081 provided by Hildenbrand et al. (2004) from Tahiti, where the 0.87 Ma collapse is followed by 1880 1082 a 90 kyr period of elevated eruptive flux, interpreted as the result of increased melt 1881 1882 1083 production. Post-collapse trace-element and isotopic signatures are consistent with higher 1883 1884 1084 degrees of mantle melting (Fig. 10). A gradual reversal of this chemical signature, 1885 1085 concomitant with a reduced eruptive flux, further supports a relationship between surface 1886 1887 32 1888 1889 1890 1891 1892 1086 loading and mantle melt fraction, although parts of this longer-term trend may be associated 1893 1087 with plate movement away from the plume head. 1894 1895 1088 1896 1089 A comparable pattern has been suggested for the evolution of Waianae volcano, Oahu, 1897 1898 1090 Hawaii, where an onshore landslide scar marks the interval between the Palehua and 1899 1900 1091 Kolekole members (Presley et al., 1997). Gradual evolution of the pre-collapse Palehua lavas 1901 1092 is consistent with a decrease in partial melting of 1-2 % through time, but the post-collapse 1902 1903 1093 Kolekole lavas mark a sharp reversal of this pattern, consistent with a 1-2 % increase in 1904 1094 partial melting (Presley et al., 1997). Kolekole lavas are less differentiated and equilibrated at 1905 1906 1095 greater depths than the preceding Palehua magmas. This pattern replicates observations at 1907 1908 1096 Fogo and La Palma, while also suggesting, like Tahiti, a relationship between growth, 1909 1097 collapse and mantle melt fraction. Presley et al. (1997) suggest that the volume of the 1910 1911 1098 Waianae slump is more than sufficient to explain the increase in partial melting implied by 1912 1099 post-collapse lava compositions. However, their assumptions may substantially overestimate 1913 1914 1100 the impacts of decompression at melt-source depths (>80 km; Watson and McKenzie, 1991) 1915 1916 1101 (cf. Manconi et al., 2009; Longpré et al., 2009). Changes in melt fraction and magma 1917 1102 compositions, comparable to those at Waianae, have not been documented for other Hawaiian 1918 1919 1103 landslides (although renewed episodes of post-collapse volcanism have been noted following 1920 1104 some of these (cf. Presley et al., 1997)), and the calculations of Manconi et al. (2009) suggest 1921 1922 1105 that, even for these very large collapses, there would not be a significant influence of 1923 1106 landslide mass redistribution on melt production. 1924 1925 1107 1926 1927 1108 9. Summary and conclusions 1928 1109 Despite severe limitations in stratigraphic reconstructions, disparate data types, and potential 1929 1930 1110 magmatic responses that span a range of timescales, a common pattern emerges from this 1931 1111 investigation of the impact of large-scale sector collapses on volcano-magmatic systems. 1932 1933 1112 Observations are consistent with an intrinsic relationship between surface loading and the 1934 1935 1113 development of mid- to upper-crustal magma reservoirs, indicating that changes to surface 1936 1114 loading can modulate magma ascent and storage, and are thereby manifested by shifts in 1937 1938 1115 erupted compositions, eruption rate and style. Discrete, rather than gradual, shifts in eruptive 1939 1116 behaviour, replicated at several volcanoes and concomitant with collapse, implies that 1940 1941 1117 wholesale and fundamental changes to the magma plumbing system can be driven by sector- 1942 1943 1118 collapse. This pattern is apparent in examples across both arc and intraplate settings. It is thus 1944 1119 reasonable to infer that the impact of large-volume sector collapse on underlying magma 1945 1946 33 1947 1948 1949 1950 1951 1120 reservoirs is independent of tectonic setting, even though data don’t currently exist to test this 1952 1121 at rift or continental intraplate volcanoes. 1953 1954 1122 1955 1123 Large sector collapses may be followed by compositionally anomalous and notably large- 1956 1957 1124 volume eruptions, which are often effusive. This short-timescale response is, however, 1958 1959 1125 dependent on the presence of eruptible (i.e. with sufficient liquid proportions) magma in the 1960 1126 crustal reservoir, and is therefore not observed in all cases. The anomalous composition of 1961 1962 1127 such eruptions implies disruption of an otherwise stable reservoir, and suggests that 1963 1128 compositions tapped from this reservoir during typical (i.e. unperturbed by collapse) periods 1964 1965 1129 of volcanism are not fully representative of crustal magma compositions. These anomalous 1966 1967 1130 events represent truly triggered eruptions (in contrast to eruption-triggered collapses, of the 1968 1131 Mount St. Helens type), and are not dependent on magma ascent driving the initial collapse. 1969 1970 1132 They thus imply that surface mass redistribution alone has the potential to disrupt stored 1971 1133 magma bodies and initiate magma ascent. Although theoretical relationships suggest that 1972 1973 1134 collapse is not necessarily expected to favour this process, such models are based on 1974 1975 1135 simplified physical and geometrical assumptions that may not well represent a vertically 1976 1136 extensive, crystal-rich plumbing system. Thus, although the absence of triggered eruptions in 1977 1978 1137 some cases could be consistent with post-collapse conditions that hinder dyke formation, they 1979 1138 could equally be explained by an absence of eruptible magma. The latter explanation is more 1980 1981 1139 consistent with subsequent compositional changes in erupted magmas at volcanoes affected 1982 1140 by large-scale sector collapse. 1983 1984 1141 1985 3-4 1986 1142 There are multiple examples of sector collapses followed by a temporary (lasting 10 years) 1987 1143 shift to eruption of more mafic compositions, often at elevated eruption rates. Such behaviour 1988 1989 1144 indicates that deeper, denser magmas can ascend to the surface without capture by a more 1990 1145 evolved upper-crustal reservoir. This implies both that surface loading may modulate mafic 1991 1992 1146 magma stalling, and that solidification of the existing upper-crustal reservoir, to the extent 1993 1994 1147 that mafic magmas can ascend to the surface, is a common result of large-scale sector 1995 1148 collapse. On longer timescales, upper-crustal storage is re-established following edifice 1996 1997 1149 regrowth, with a transition towards eruption of more evolved compositions at eruption rates 1998 1150 comparable to pre-collapse activity. Volcano growth and collapse, with a co-developing 1999 2000 1151 crustal plumbing system, can therefore be defined within a broadly cyclic pattern of 2001 4-5 2002 1152 behaviour, on timescales of 10 years in arc settings. This supports theoretical relationships 2003 1153 that have previously been proposed, although stratigraphic reconstructions remain too coarse 2004 2005 34 2006 2007 2008 2009 2010 1154 and incomplete to provide a quantitative analysis of this cyclicity. Broadly similar behaviour 2011 1155 is observed in intraplate ocean-island settings, but the absence of an evolved shallow 2012 2013 1156 reservoir in many instances results in less clear compositional shifts than those outlined 2014 1157 above. Nevertheless, post-collapse changes in behaviour at ocean islands can be explained via 2015 2016 1158 plumbing system disruption, and enhanced melt production is not a necessary part of the 2017 2018 1159 general relationship between collapse and subsequent eruptive behaviour. 2019 1160 2020 2021 1161 The relationships described above are evident in multiple examples of large-volume sector 2022 1162 collapses, generally with volumes exceeding 5 km 3. These events are significantly larger than 2023 2024 1163 historical sector collapses, with the exception of Ritter Island. The absence of such clear post- 2025 2026 1164 collapse responses following historical examples may thus be attributed to the smaller scale 2027 1165 of these events, both in absolute terms, and potentially as a proportion of the reservoir and 2028 2029 1166 edifice volumes. The load redistribution associated with smaller collapses may be comparable 2030 1167 to volumetric and pressure changes during typical eruptive behaviour, and insufficient to 2031 2032 1168 drive major changes to the plumbing system (the impacts may potentially be accommodated 2033 2034 1169 by decompression within the reservoir (e.g., Voight et al., 2010), and a steady-state is thus 2035 1170 maintained). In contrast, the largest sector collapses, much less frequent in the history of 2036 2037 1171 individual volcanoes, result in disequilibrium between the crustal reservoir and surface load, 2038 1172 marking major shifts in the long-term development of a volcano-magmatic system. The 2039 2040 1173 magnitude of collapse is thus significant in terms of its magmatic impact, although 2041 1174 investigating the detail of this relationship is again constrained by limitations in 2042 2043 1175 reconstructing collapse volumes and magma fluxes. 2044 2045 1176 2046 1177 It is notable that surface mass redistribution alone can drive reorganisation of the upper 2047 2048 1178 crustal plumbing system, without large scale magmatic eruption. In such cases, magma 2049 1179 within the pre-collapse reservoir remains in-situ, yet post-collapse mafic activity indicates 2050 2051 1180 that subsequent eruptions can be fed by magma ascending through this reservoir. 2052 2053 1181 Furthermore, isotopic, mineralogical and trace-element characteristics of post-collapse 2054 1182 evolved magmas imply that post-collapse upper-crustal reservoirs, although occupying a 2055 2056 1183 similar depth range, represent a discrete plumbing system, not directly related to the pre- 2057 1184 collapse storage system. Taken together, these observations are consistent with a vertically 2058 2059 1185 extensive and crystal-dominated crustal plumbing system beneath volcanoes, since it is 2060 2061 1186 difficult to envisage a large-volume, liquid dominated system resulting in such sharp changes 2062 1187 in behaviour. 2063 2064 35 2065 2066 2067 2068 2069 1188 2070 1189 The patterns discussed here are based on a relatively small number of examples, representing 2071 2072 1190 all available cases where chemical, chronological, stratigraphic and volumetric data are 2073 1191 sufficient to assess pre- and post-collapse patterns in volcanism. Even in these instances, the 2074 2075 1192 impact of collapse can only be analysed at a low resolution. Further testing of the processes 2076 2077 1193 outlined here, and the theoretical relationships that imply episodic development of volcanic 2078 1194 systems and modulation of magma storage by surface loading, is reliant on a greater number 2079 2080 1195 of more detailed investigations of individual volcanoes affected by sector collapse, 2081 1196 combining compositional and stratigraphic information on timescales spanning tens of 2082 2083 1197 thousands of years. The value of such studies lies not simply in improving the understanding 2084 2085 1198 of individual systems, but in providing essential insights into the fundamental controls on 2086 1199 magma storage, eruptive behaviour, and the nature of magma reservoirs. 2087 2088 1200 2089 1201 Acknowledgements 2090 1202 This work was supported by Natural Environment Research Council (NERC) grant 2091 1203 NE/I02044X/1 & 2. Discussions with V. 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Geology, 32: 573-576. 2757 1769 Wörner, G., Harmon, R.S., Davidson, J., Moorbath, S., Turner, D.L., McMillan, N., Nye, C., 2758 1770 Lopez-Escobar, L., Moreno, H., 1988. The Nevados de Payachata volcanic region 2759 1771 (18°S/69°W, N. Chile). Bulletin of Volcanology, 50: 287-303. 2760 1772 Yamamoto, T., Nakamura, Y., Glicken, H., 1999. Pyroclastic density current from the 1888 2761 1773 phreatic eruption of Bandai volcano, NE Japan. Journal of Volcanology and 2762 1774 Geothermal Research, 90: 191-207. 2763 1775 Yoshida, H., 2013. Decrease of size of hummocks with downstream distance in the rockslide- 2764 1776 debris avalanche deposit at Iriga volcano, Philippines: similarities with Japanese 2765 1777 avalanches. Landslides, 10: 665-672. 2766 1778 Zellmer, G.F., Hawkesworth, C.J., Sparks, R.S.J., Thomas, L.E., Harford, C.L., Brewer, T.S., 2767 2768 1779 Loughlin, S.C., 2003. Geochemical evolution of the Soufriere Hills volcano, 2769 1780 Montserrat, Lesser Antilles volcanic arc. Journal of Petrology, 44: 1349-1374. 2770 1781 Zernack, A.V., Cronin, S.J., Bebbington, M.S., Price, R.C., Smith, I.E.M., Stewart, R.B., 2771 2772 47 2773 2774 2775 2776 2777 1782 Procter, J.N., 2012. Forecasting catastrophic stratovolcano collapse: A model based 2778 1783 on Mount Taranaki, New Zealand. Geology, 40: 983-986. 2779 1784 Zernack, A.V., Cronin, S.J., Neall, V.E., Procter, J.N., 2011. A medial to distal volcaniclastic 2780 1785 record of an andesite stratovolcano: detailed stratigraphy of the ring-plain succession 2781 1786 of south-west Taranaki, New Zealand. International Journal of Earth Sciences, 100: 2782 1787 1937-1966. 2783 1788 Zernack, A.V., Procter, J.N., Cronin, S.J., 2009. Sedimentary signatures of cyclic growth and 2784 1789 destruction of stratovolcanoes: a case study from Mt. Taranaki, New Zealand. 2785 1790 Sedimentary Geology, 220: 288-305. 2786 1791 2787 2788 2789 2790 2791 2792 2793 2794 2795 2796 2797 2798 2799 2800 2801 2802 2803 2804 2805 2806 2807 2808 2809 2810 2811 2812 2813 2814 2815 2816 2817 2818 2819 2820 2821 2822 2823 2824 2825 2826 2827 2828 2829 2830 2831 48 2832 2833 2834 2835 a 2836 1792 Table 1 : Theoretical impacts of sector collapse on an upper crustal liquid magma body . 2837 Pressure parameters ( P) (Fig. 3) b and impact Sector collapse type 2838 Pr(f) < P r(i) Pr(f) > P r(i) 2839 Eruption- if P

P c(f) m(f) if P

P r(f) otherwise, 2847 shorter time to longer time to eruption, lower 2848 triggered eruption, rate 2849 eruption higher rate 2850 No eruptible magma when eruptible magma forms, when eruptible magma forms, 2851 eruption favoured relative to eruption impeded relative to 2852 pre-collapse conditions pre-collapse conditions 2853 a 2854 cf. Pinel and Albino (2013) 2855 b Subscript (i) refers to pre-collapse conditions; (f) refers to post-collapse conditions 2856 1793 2857 1794 Table 2 : Summary of historical sector collapses (volumes >1 km 3)a. 2858 Location Date Collapse Preceding repose Precursory Post-collapse activity 2859 (A.D.) volume interval activity 2860 (km 3) 2861 Oshima- 1741 2.5b ~1500 years ~10 days of Strong explosive eruption 2862 Oshima, Japan explosive eruptions accompanying collapse; ~9 months 2863 intense activity; minor sporadic 2864 activity to 1790; no subsequent eruption 2865 Ritter Island, 1888 4.2; 2.4d 1-3 years?; frequent Uncertain Compositionally bimodal post- 2866 Papua New minor explosive collapse eruption, including an 2867 Guinea c eruptions evolved pumiceous component not 2868 known from any other Ritter 2869 samples. Extent of eruption- 2870 associated turbidite, immediately 2871 overlying collapse deposits, 2872 suggests a powerful submarine explosive eruption triggered by 2873 collapse. Smaller subsequent 2874 eruptions have built a submarine 2875 scoria cone within the scar, 2876 compositionally distinct from pre- 2877 collapse samples and involving 2878 both more mafic and more evolved 2879 components. e 2880 Bandai, Japan 1888 0.5 No magmatic 1 week of minor Immediate strong phreatic eruption. activity for ~25 seismicity; event No subsequent activity. 2881 kyr; phreatic triggered by M5 2882 explosions in 1808 earthquake 2883 Shiveluch, 1964 1.15 Preceding eruption Seismicity over 10 Immediate phreatic explosion; 2884 Kamchatka in 1946-50; months, Plinian eruption within minutes, but 2885 frequent prior accelerating before no directed blast. Broadly 2886 activity (effusive collapse continuous extrusion with 2887 and explosive) explosions from 1980 to present. 2888 Mount St. 1980 2.3-2.5 Last eruption 2 months of Immediate lateral blast and Plinian 2889 2890 49 2891 2892 2893 2894 Helens, U.S.A. ~1857; frequent increasing explosive eruption, followed by 2895 prior activity seismicity, slope lava-dome extrusion; subsequent 2896 deformation and phases of dome extrusion. 2897 phreatic explosions. 2898 a Data from Global Volcanism Program (2013) and: Oshima-Oshima (Katsui and Yamamoto, 1981; Satake and Kato, 2899 2001; Satake, 2007); Ritter (Johnson, 1987; Day et al., 2015; Karstens et al., 2019; Watt et al., 2019); Bandai 2900 (Yamamoto et al., 1999; Yoshida, 2013); Shiveluch (Belousov, 1995; Ponomareva et al., 2006); Mount St. Helens 2901 (Voight et al., 1983; Glicken, 1996). b 3 2902 The subaerial collapse scar suggests a much smaller volume (0.4 km ); the larger volume is based on relatively limited bathymetry, but is consistent with tsunami observations (Satake and Kato, 2001). 2903 c The base of the scar is several hundred metres below sea-level, meaning that replacement by seawater partially 2904 compensates for the mass-removal associated with collapse. 2905 d Volumes based on reconstructions in Day et al. (2015) and Karstens et al. (2019) respectively. 2906 e An earlier quoted volume of 1.4 km 3 is disputed by later authors, whose topographic reconstructions suggest a 2907 volume of ≤0.5 km3 (cf. Yoshida, 2013). 2908 2909 2910 2911 2912 2913 2914 2915 2916 2917 2918 2919 2920 2921 2922 2923 2924 2925 2926 2927 2928 2929 2930 2931 2932 2933 2934 2935 2936 2937 2938 2939 2940 2941 2942 2943 2944 2945 2946 2947 2948 2949 50 2950 2951 2952 2953 a 2954 Table 3 : Sector collapses from arc volcanoes associated with an identified change in magma composition or eruptive behaviour . 2955 Volcano Collapse Age (ka) Volume Observations: chemical and petrological Observations: eruption rate and style Sources name (km 3) 2956 Shiveluch, Old 16-10 ~30 A greater predominance of more evolved Frequent large explosive eruptions over Belousov et al. 2957 Kamchatka Shiveluch (intermediate) magmas post-collapse. Evolved 3kyr post collapse, then infrequent. Pre- 1999; Gorbach 2958 compositions are broadly similar pre- and post- collapse edifice very large but et al., 2013 2959 collapse, but mafic compositions are distinct (pre- structurally stable; post-collapse marked 2960 collapse have lower Mg#, higher Fe, Ti, and distinct by frequent repetitive small collapses. 2961 trace element compositions (e.g. Cr, V)). Increases in magma viscosity post- 2962 collapse; a dominance of lava domes over 2963 flows, leading to a steeper and possibly 2964 less stable edifice. Nevado de 18.5 6-12 Collapse deposit overlain by PDC deposits with Robin et al., 2965 Colima, mixed juvenile material, including the most mafic 1990; Stoopes 2966 Mexico products known from the volcano. Petrological and Sheridan, 2967 observations (reaction rims, resorption) and textural 1992; Capra et 2968 evidence of mixing in juvenile material indicate al., 2002 2969 mixing of mafic and felsic magmas prior to eruption. 2970 Later lavas (basaltic andesites) are compositionally 2971 similar to those preceding collapse. 2972 Fuego de early- to ~10 Collapse preceded by dacites, but subsequent PDCs Thick andesite lava flows precede Robin et al., Colima, mid- contain basaltic andesite juvenile material, and collapse; explosive andesitic eruptions, at 1990; Stoopes 2973 Mexico Holocene andesites dominate subsequent cone rebuilding. higher frequency for ~1 kyr after and Sheridan, 2974 collapse, have dominated subsequent 1992; Capra et 2975 cone building. al., 2002 2976 Orizaba, Torrecillas/ 210 ~25 Poorly resolved shift in dominant mineral Carrasco-Núñez 2977 Mexico Jamapa assemblage from pyroxene andesite to hornblende et al., 2006; 2978 andesite in pre- and post-collapse units. Schaaf and 2979 Carrasco-Núñez, 2980 2010 Soufrière Deposit 130 6-10 Sharp departure from pre-collapse two-pyroxene Available dates suggest elevated eruption Watt et al., 2981 Hills, 2/South andesites to subsequent basaltic volcanism in the rates and a duration of less than a few 2012; Cassidy et 2982 Montserrat Soufrière South Soufrière Hills episode. Younger andesites are thousand years for the South Soufrière al., 2015a 2983 Hills petrographically distinct (hornblende-phyric) from Hills episode. 2984 pre-collapse andesites. 2985 Conil/Pelée, Le Prêcheur, 126 15 Collapse closely bracketed by two extremely similar Germa et al., 2986 Martinique D1 lavas (possible extrusion-driven collapse, or post- 2011, 2015; 2987 collapse eruption driven by reservoir unloading). Boudon et al., 2988 2989 51 2990 2991 2992 2993 2994 Collapse marks the end of the Conil complex 2013 2995 (andesitic); subsequent early stages of Pelee 2996 complex dominated by basaltic andesite. 2997 Pelée, St Pierre, D2 ~32 9 Collapse preceded by acid andesites and followed by Increased eruption frequency in 5 kyr Boudon et al., 2998 Martinique basaltic andesite eruptions from ~32-27 ka, the post-collapse period (based on marine 2013; Germa et 2999 system then returning to acid andesite eruptions tephra deposits) and change in al., 2015 3000 from ~22.5 ka. predominant style from dome-forming to 3001 explosive behaviour. 3002 Tungurahua, Tungurahua 3 6 Collapse of the andesitic TII edifice was followed by Post-collapse dacite lavas are unusually Hall et al., 1999 3003 Ecuador II extrusion of mineralogically unusual dacite lavas (ol, extensive (6 km flow lengths). cpx, opx and amph-phyric) before a 700-yr pause in Subsequent cone rebuilding activity at 3004 volcanism, followed by eruption of homogeneous high eruptive flux (1.5 km 3/kyr). 3005 basaltic andesites (2.3-1.4 ka), distinct from pre- 3006 collapse magma compositions. Eruptions since 1.2 3007 ka include more evolved compositions. Collapse of 3008 the TI replicates this broad pattern. 3009 Chimborazo, Riobamba 65-60 10-12 Collapse marks a permanent shift in magma Voluminous and homogeneous andesitic Samaniego et 3 3010 Ecuador chemistry; post-collapse rocks are more mafic, lack lava flows (1-1.5km ) erupted post- al., 2012 3011 amphibole as a phenocryst phase, and are enriched collapse. Large volume of flows notable in Mg, K and some trace elements (e.g. Rb, Th) in long term history of volcano. 3012 relative to pre-collapse. 3013 Tata Sabaya, ~6 6 Post-collapse thick lava domes and flows, all of very de Silva et al., 3014 Bolivia similar composition. Similar compositions 1993; Godoy et 3015 throughout history of volcano, but two possibly early al., 2012 3016 post-collapse domes contain inclusions more mafic 3017 than any other known products at the volcano. Other 3018 post-collapse lavas are more porphyritic and more 3019 evolved, but the relative timing of post-collapse eruptions is poorly constrained. 3020 Parinacota, Parinacota 9 6 Post-collapse magmas more mafic than pre-collapse Very high eruptive flux (up to 10 km 3/ Wörner et al., 3021 Chile rocks and include the most mafic products known at kyr) for up to two thousand years post- 1988; Hora et 3022 the volcano (from flank vents), although these fall collapse. Rapid regrowth has buried al., 2007, 2009; 3023 on a longer term trend to more mafic compositions. collapse scar. Jicha et al., 2015 3024 Post-collapse magmas geochemically consistent with 3025 rapid crustal transit and limited upper-crustal 3026 storage, in contrast with pre-collapse rocks. 3027 Phenocryst assemblage changes from hornblende- andesite to post-collapse two-pyroxene (Hbl absent) 3028 3029 3030 52 3031 3032 3033 3034 3035 assemblages. 3036 Planchón, Rio Teno 11 4-5 Basaltic rocks predominate at Planchón, but these Post-collapse vent migrated to focus Tormey et al., 3037 Chile evolved pre-collapse to basaltic andesite, with within centre of collapse scar (cf. 1995; Tormey, 3038 evidence for upper crustal storage. Post-collapse Stromboli). Higher proportion of 2010 3039 compositions revert to more mafic basalts, similar to pyroclastic material (i.e. explosive 3040 earlier Planchón lavas. activity) relative to pre-collapse volcano. 3041 San Pedro mid- ~4 Collapse was followed by eruption of ~1 km 3 The composite post-collapse flow is Costa and 3042 (Tatara – Holocene compositionally variable silicic lavas, containing relatively voluminous, with a total Singer, 2002 3 3043 San Pedro), basaltic and gabbroic inclusions, and an associated volume of 0.8 km contrasting with 0.2 3 3044 Chile explosive deposit. These lavas are interpreted as km total volume for subsequent representing eruption from a zoned storage system; Holocene summit flows. 3045 they are petrologically (e.g. amphibole-phyric) and 3046 chemically (Sr isotope and trace element signatures) 3047 distinctive from pre-collapse basaltic andesite and 3048 andesites, and younger basaltic andesites, which are 3049 also distinct from each other. 3050 Antuco, 6 5 Pre-collapse rocks span basalts to dacites; the Debris avalanche overlain by two Lohmar et al., 3051 Chile earliest post-collapse lavas are basaltic andesite, but unusually voluminous lava flows, 2005; Martínez 3052 followed by more mafic explosive and effusive compositionally identical, and distinct et al., 2018; eruptions that have infilled much of the collapse scar from the more mafic units that have Watt, 3053 and persisted for the rest of the Holocene. dominated subsequent activity, and have unpublished data 3054 been erupted at a higher flux than pre- 3055 collapse magmas. 3056 Stromboli, Upper 14 1.4-3.1 Pre-collapse latite to trachyte compositions are Effusive eruptions dominate the Francalanci et 3057 Italy Vancori followed by relatively monotonous highly Neostromboli period (13-6 ka) at a higher al., 1989; 3058 porphyritic shoshonites, the most potassic rocks eruptive flux than the pre-collapse period. Hornig- 3059 known at the volcano and much more basic than any Kjarsgaard et al., 3060 rocks from the preceding period. More evolved 1993; Tibaldi, biotite shoshonites, consistent with fractional 2001, 2004; 3061 crystallisation of preceding magmas, erupt towards Romagnoli et 3062 the end of the period (6 ka), implying re- al., 2009; 3063 establishment of upper crustal storage. Vezzoli et al., 3064 2014 3065 Stromboli, Neostromboli 6 ~1 The collapse-associated Secche di Lazzaro deposits Collapse associated Secche di Lazzaro Bertagnini and 3066 Italy share petrographic and geochemical characteristics deposits represent the most intense Landi, 1996; 3067 with late-Neostromboli evolved shoshonites, but are explosive activity of the past 6 kyr, Tibaldi, 2001; 3068 mineralogically and compositionally heterogeneous, suggestive of collapse-driven Petrone et al., consistent with decompression (i.e. collapse) driven decompression of the shallow plumbing 2009; Vezzoli et 3069 3070 3071 53 3072 3073 3074 3075 3076 disruption of the shallow reservoir. No evidence of and hydrothermal system. Younger al., 2014 3077 fresh mafic input prior to collapse. Subsequent activity dominated by strombolian 3078 activity shifts to basaltic-shoshonitic magmas, behaviour. 3079 continuing across younger small sector collapses. 3080 Ritter AD 1888 2.4-4.2 See Table 2. See Table 2. Day et al., 2015; 3081 Island, Karstens et al., 3082 Papua New 2019; Watt et 3083 Guinea al., 2019 a 3084 Additional examples with more limited data include Santa Ana, El Salvador (Siebert et al., 2004), Ollagüe, Chile/Bolivia (Feeley et al., 1993), St Lucia and Martinique 3085 (Boudon et al., 2013) and Guadeloupe (Samper et al., 2007), as discussed in the text. 3086 3087 3088 3089 3090 3091 3092 3093 3094 3095 3096 3097 3098 3099 3100 3101 3102 3103 3104 3105 3106 3107 3108 3109 3110 3111 3112 54 3113 3114 3115 3116 3117 3118 Table 4 : Sector collapses from intraplate ocean-island volcanoes associated with an identified change in magma composition or eruptive 3119 behaviour. 3120 Island / Collapse Age (ka) Volume Observations: chemical and petrological Observations: eruption rate and style Sources volcano name (km 3) 3121 Tenerife, Icod 175 80-150 Large-volume phonolitic explosive eruptions Mafic lavas emplaced at elevated rate Ablay and 3122 Canary preceded event, but may also have accompanied or over 10 kyr post-collapse period, with Hürlimann, 3123 Islands a been triggered by latter stages of failure (Abrigo broadly exponential subsequent decline in 2000; Masson et 3124 ignimbrite). Subsequent lavas infilling the base of the output rates associated with increasing al., 2002; 3125 scar are mafic, transitioning to more evolved prevalence of more evolved compositions. Carracedo et al., 3126 compositions over tens of kyr, with eventual 2007; Boulesteix 3127 establishment of shallow magma storage and growth et al., 2012; 3128 the phonolitic Teide complex. Paris et al., 3129 2017; Hunt et al., 2018 3130 Tenerife, Micheque 830 Early post-collapse rocks are mafic ankaramites and Infilling of the collapse scar suggests high Carracedo et al., 3131 Canary plagioclase basalts. Later eruptions of intermediate eruption rates of post-collapse volcanism 2011 3132 Islands b and evolved (trachytes) magmas, within the collapse (particularly in the initial stages), focused 3133 scar, suggest that centralised volcanism within the within the scar and promoting shallow 3134 scar promoted shallow magma storage and magma storage and differentiation. 3135 differentiation, across a 300 kyr period. 3136 Tenerife, Teno 6100- >20-25 The youngest pre-collapse rocks include the most Lapilli-tuffs in post-collapse breccia Longpré et al., 3137 Canary massif, two 5900 each silicic and least magnesian lavas from the shield sequences suggest explosive eruptions 2009 Islands unnamed stratigraphy, while rocks above the collapse following collapse (derived from the 3138 landslides unconformities are notably mafic, marking a sharp disrupted shallow plumbing system), 3139 transition. The pattern is repeated for each landslide. anomalous in the lava-dominated shield. 3140 Dense, crystal-rich ankaramites are also notably 3141 frequent in the immediate post-collapse sequences. 3142 La Palma, Cumbre 560 95 Post-collapse growth of Bejenado volcano across a Elevated growth rates in early post- Guillou et al., 3143 Canary Nueva shallowing trend (~25-18 km) and showing evidence collapse stages. 1998; Masson et 3144 Islands of increasing differentiation and reduced magma al., 2002; Galipp 3145 supply rates through time. Storage depths distinct et al., 2006 from pre-collapse magmas, suggesting development 3146 of a distinct post-collapse plumbing system. 3147 El Hierro, El Golfo 87-39 247 Pre-collapse units include Mg-poor trachytes (~60 Manconi et al., 3148 Canary wt% SiO 2), but evolved compositions are absent in 2009; Longpré 3149 Islands the post-collapse stratigraphy, dominated by dense, et al., 2011; 3150 crystal-rich and high-MgO basanites (mean of 44 León et al., 2017 3151 wt% SiO 2). 3152 3153 55 3154 3155 3156 3157 3158 Fogo, Cape Monte 123-62 130-160 Early post-collapse magmas stored at lower depths Post-collapse vent migration is consistent Day et al., 3159 Verde Amarelo than immediately preceding magmas (although all with modified magma ascent pathways 1999b; Amelung 3160 depths are uppermost mantle, with no evidence for and development of a distinct plumbing and Day, 2002; 3161 prolonged crustal storage). Later eruptions span system. Hildner et al., 3162 broader and shallower pressure ranges, suggesting 2012; 3163 increasing complexity of plumbing system Maccaferri et al., 3164 established upon regrowth. 2017 3165 Tahiti N. landslide 870 >500 Sharp shifts in La/Yb, and Sr and Nd isotope ratios in 90 kyr post-collapse period of elevated Hildenbrand et 3 3166 post-collapse magmas are consistent with increased output (~5 km /kyr), interpreted as arising al., 2004 decompression melting, with further isotopic from increased mantle melt fraction. 3167 evolution and transition to silica-undersaturated Gradual reduction in productivity and 3168 compositions suggesting reduced melting upon output rate upon regrowth. 3169 regrowth. 3170 Volcan <100 unknown A high proportion of post-collapse magmas with Post-collapse shift in dominant vent Geist et al., 2002 3171 Ecuador, elevated Mg# and in some cases lying on distinct locations, suggesting either bypassing of 3172 Isabela, trace-element trends (shifted to higher Mg#) relative formerly dominant shallow reservoir, or 3173 Galapagos to the more consistent compositions of pre-collapse destruction of this reservoir following 3174 lavas. collapse. 3175 Waianae, Waianae 2980 up to Collapse separates Palehua (older) and Kolekole Presley et al., Oahu, slump 6100 (younger) members, with Palehua showing temporal 1997 3176 Hawaii c evolution towards lower degrees of partial melting 3177 (based on Nb/Zr), and the post-collapse onset of 3178 Kolekole sharply reversing this trend, being 3179 comparable to early Palehua lavas. Higher Mg 3180 contents and bulk chemistry suggest Kolekole lavas 3181 evolved at greater depths than Palehua, fractionating 3182 ol + cpx, with lower overall extents of fractionation than Palehua. 3183 a An older (but potentially much smaller volume) collapse to the south of the caldera may have been eruption triggered (Dávila Harris et al., 2011; Paris et al., 2017). 3184 b Other collapses on N.E. rift zone (Orotava, Güímar) occurred at a similar time and after the Micheque collapse, but the strongest evidence of a post-collapse control on 3185 subsequent rift volcanism is within the infilled Micheque collapse scar (Carracedo et al., 2011). . 3186 c More limited evidence of post-collapse rejuvenated volcanism exists for other Hawaiian volcanoes (cf. Presley et al., 1997), while post-collapse plumbing system 3187 modification, and possible chemical changes and an episode of phreatomagmatic activity, were associated with the SW Mauna Loa landslides (Lipman et al., 1990). 3188 3189 3190 3191 3192 3193 3194 56 3195 3196 3197 3198 3199 3200 1796 Figure 1 3201 1797 Example morphologies of volcanoes affected by sector collapse. A: Oblique view of Mount 3202 1798 St Helens, USA (image: NASA Earth Observatory), showing the approximate distribution of 3203 1799 the landslide mass (Glicken, 1996) following the 1980 sector collapse. The distance across 3204 1800 the collapse amphitheatre is 2 km. The deposit volume of 2.5 km 3 was distributed up to 29 3205 1801 km from the volcano (Glicken, 1996). B: Gradient-shaded bathymetry of Ritter Island, Papua 3206 1802 New Guinea (image: Christoph Böttner), site of the largest historical sector collapse, in 1888, 3207 1803 showing the submerged island morphology (the shallowest edifice and subaerial remnant is 3208 1804 not shown). The collapse was approximately twice the volume of that at Mount St. Helens 3209 1805 (Day et al., 2015). C: Scale cross-sections (without vertical exaggeration) through Socompa, 3210 1806 Chile (one of the largest known subaerial sector collapses; Wadge et al., 1995) and Ritter 3211 1807 Island (Ward and Day, 2003), both composite volcanoes in arc settings, and El Hierro, 3212 1808 Canary Islands (Masson et al., 2002), one of the youngest examples of a large-scale flank 3213 1809 landslide on an oceanic intraplate volcanic island. The pre-failure surfaces are conjectural. 3214 3215 1810 3216 1811 Figure 2 3 3217 1812 A summary of documented sector collapses (>1 km ). A: Global distribution of known 3218 1813 examples, divided by tectonic setting. B: Relative proportions of sector collapse volumes 3219 1814 across different tectonic settings (for events where specific volume estimates are available). 3220 1815 Most subduction-zone collapses are <5 km 3, and the very largest collapses are all from 3221 1816 intraplate ocean islands. 3222 1817 3223 1818 Figure 3 3224 1819 A theoretical framework for investigating the effects of edifice loading on magma chamber 3225 1820 pressure, dyke formation and eruption rate (cf. Pinel et al., 2010). A: An edifice load (ℎ ) 3226 1821 affects the stress field around an upper crustal liquid cavity, characterised by a pressure Pm, 3227 1822 and connected to a deeper source with pressure P and density • . P describes the critical 3228 S r 3229 1823 pressure for magma chamber rupture and dyke formation, and Pc the pressure at which this 3230 1824 dyke will close. Any change in edifice load (depicted in red) will alter Pm, Pr and Pc. B: The 3231 1825 rate of pressurisation of a shallow cavity connected to a deeper source depends on the 3232 1826 pressure difference between the two (shown by the black curve), predicting a constant 3233 1827 eruption rate under any particular set of conditions. If Pr is relatively low (condition 1), then 3234 1828 replenishment of the upper chamber (following the pressure drop to Pc) will be relatively 3235 1829 rapid. At higher Pr (condition 2), a longer pressurisation time results in a lower eruption rate. 3236 1830 C: As an edifice load increases, Pr falls to a minimum and then increases with further edifice 3237 1831 loading, with the eruption rate therefore following an opposite pattern. Sector collapse results 3238 1832 in an instantaneous reduction in edifice load. This may either favour (condition 1) or impede 3239 1833 (condition 2) subsequent eruptions, depending on the impact of unloading on Pr. 3240 1834 3241 1835 Figure 4 3242 1836 Long-term eruptive flux at well studied composite arc volcanoes. The main graph shows 3243 1837 estimates of eruptive flux through time (<400 ka) compiled from detailed field studies at 3244 3245 1838 thirteen volcanoes (Hildreth and Lanphere, 1994; Singer et al., 1997, 2008; Hobden et al., 3246 1839 1999; Davidson and de Silva, 2000; Thouret et al., 2001; Hildreth et al., 2003a, 2003b; Frey 3247 1840 et al., 2004; Bacon and Lanphere, 2006; Jicha and Singer, 2006; Hora et al., 2007, 2009; 3248 1841 Samaniego et al., 2012). Relatively short episodes of heightened output are often interspersed 3249 1842 with longer, quieter periods. The average of all datasets is shown as a red line (with shaded 3250 1843 upper and lower quartiles). The inset graph shows literature estimates of eruptive flux against 3251 1844 the duration of the estimate (data from Wadge, 1982; Hall et al., 1999; Thouret et al., 2001, 3252 3253 3254 57 3255 3256 3257 3258 3259 1845 2005; White et al., 2006; Singer et al., 2008; Samaniego et al., 2012; Zernack et al., 2012), 3260 1846 showing that decreasing temporal resolution leads to lower apparent flux. 3261 1847 3262 1848 Figure 5 3263 1849 Examples of anomalous post-collapse eruptions, interpreted as representing collapse-driven 3264 1850 disruption of a pre-existing magma reservoir. A: Post-collapse lava flows at Antuco, Chile, 3265 1851 directly overlying the mid-Holocene sector collapse deposit. The two lavas have near 3266 1852 identical compositions, and are less mafic than the basalts that have dominated subsequent 3267 1853 cone rebuilding in the collapse scar. These lavas are unusually voluminous and extensive in 3268 1854 the context of both older and younger activity. B: Compositionally anomalous post-collapse 3269 1855 products at Nevado de Colima, Mexico (Robin et al., 1987, 1990) and San Pedro, Chile 3270 1856 (Costa and Singer, 2002). In both cases debris avalanche deposits are directly overlain by 3271 1857 units that span a wide bulk compositional range, and in the case of San Pedro also have a 3272 1858 distinctive phenocryst assemblage relative to both pre- and post-collapse products. Post- 3273 collapse lavas at Tungurahua, Ecuador, provide a comparable example (Hall et al., 1999). 3274 1859 3275 1860 3276 1861 Figure 6 3277 1862 Post-collapse shifts to more mafic magma compositions (highlighted by bulk MgO and SiO 2 3278 1863 compositions) for collapses at Soufrière Hills, Montserrat (Zellmer et al., 2003; Cassidy et al., 3279 1864 2012, 2015a), Pelée, Martinique (Boudon et al., 2013), Stromboli, Italy (Hornig-Kjarsgaard et 3280 1865 al., 1993; Vezzoli et al., 2014), and Parinacota, Chile (Hora et al., 2009). Data are plotted at 3281 1866 the same timescale and demonstrate the data gaps and differences in temporal resolution that 3282 1867 hinder inter-volcano comparisons and reconstructions of eruptive behaviour associated with 3283 1868 prehistoric sector collapses. When specific ages are unavailable, the stratigraphic intervals of 3284 1869 data are indicated by vertical arrows. Soufrière Hills and Pelée returned to eruption of more 3285 1870 evolved compositions, comparable but not identical to pre-collapse rocks. Further examples 3286 1871 of more mafic magmatism in post-collapse cone-building episodes are provided in Table 3. 3287 1872 3288 1873 Figure 7 3289 1874 Persistent shifts in bulk magma compositions following sector collapse, at Shiveluch, 3290 1875 Kamchatka (Gorbach and Portnyagin, 2011; Gorbach et al., 2013) and Chimborazo, Ecuador 3291 3292 1876 (Samaniego et al., 2012). The sharp nature of these shifts, replicated at several other 3293 1877 volcanoes, are interpreted as indicating that the post-collapse reservoir is discrete 3294 1878 (compositionally, geometrically or thermally) from the preceding system, with resultant 3295 1879 changes in the relative crustal influences on erupted magma compositions. 3296 1880 3297 1881 Figure 8 3298 1882 A conceptual model of the impact of large-volume sector collapse at an arc volcano. The top 3299 1883 panel depicts the pre-collapse state, with a magma reservoir in equilibrium with the surface 3300 1884 load. Short term post-collapse behaviour is dependent on the presence of eruptible magma, 3301 1885 which may be destabilised by the surface mass redistribution. Cone rebuilding is promoted by 3302 1886 subsequent ascent of deeper mafic magmas, which may reach the surface with little 3303 1887 modification. Longer-term regrowth promotes upper crustal storage and a return to more 3304 1888 evolved compositions, but with a plumbing system that is spatially and temporally distinct 3305 1889 from the pre-collapse reservoir. 3306 1890 3307 1891 Figure 9 3308 3309 1892 Sector collapse and magmatic processes on Tenerife, Canary Islands. The map shows three 3310 1893 morphologically prominent collapse scars, the youngest of which is the Icod landslide (~175 3311 1894 ka; Boulesteix et al., 2012), as well as the location of the Micheque collapse scar (Carracedo 3312 3313 58 3314 3315 3316 3317 3318 1895 et al., 2011). Mafic vent sites (from Dóniz-Páez, 2015) highlight the rift zone structure of the 3319 1896 island. More evolved magmas only erupt in the central region (Ablay et al., 1998), which is 3320 1897 also the location of the Cañadas caldera structure, and inferred to mark a long-lived upper 3321 1898 crustal magma reservoir. The central panel shows reconstructed eruption rates in the caldera 3322 1899 region following the Icod collapse, and the right panel shows bulk rock MgO and K 2O 3323 1900 contents for the same area (based on extrapolated ages, and both derived from the lava 3324 1901 stratigraphy in Boulesteix et al. (2012)). This highlights a sharp-shift to mafic compositions 3325 1902 following the Abrigo ignimbrite eruptions (composition from Wolff et al. (2000)), 3326 1903 temporarily elevated and then declining eruption rates, and a trend towards more evolved 3327 1904 compositions as upper-crustal magma storage is re-established over a ~150 kyr period. 3328 1905 3329 1906 Figure 10 3330 1907 Long-term magmatic trends at various ocean islands indicative of a direct collapse-driven 3331 1908 influence on magma plumbing systems. The left panel shows calculated storage depths (from 3332 clinopyroxene-melt barometry) for eruptive units from La Palma, Canary Islands (Galipp et 3333 1909 3334 1910 al., 2006) and Fogo, Cape Verde Islands (Hildner et al., 2012), highlighting gradual 3335 1911 shallowing trends over ~100 kyr time periods, with a shift to higher pressures following 3336 1912 sector collapse. The right panel shows changes in trace element chemistry of lavas on Tahiti 3337 1913 (Hildenbrand et al., 2004, plotting only lavas with a differentiation index <30), either side of 3338 1914 the major ~870 ka collapse. A compositional shift coincides closely with collapse, and 3339 1915 suggests an interruption of long-term evolutionary trends. This is interpreted by Hildenbrand 3340 1916 et al. (2004) as reflecting a relationship between loading, collapse, and mantle partial melting, 3341 1917 with a collapse-driven increase in melt fraction. 3342 1918 3343 3344 3345 3346 3347 3348 3349 3350 3351 3352 3353 3354 3355 3356 3357 3358 3359 3360 3361 3362 3363 3364 3365 3366 3367 3368 3369 3370 3371 3372 59 3373       

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