Calderas are among the most active and dangerous volcanoes. Caldera unrest is defined by enhanced seismicity, gravity changes, surface deformation, and degassing. Although much caldera unrest does not lead to an eruption, every eruption is preceded by an unrest episode. Therefore, the proper description of unrest and the forecast of its possible outcome is a timely and challenging task. Here we review the best known unrest at calderas from 1988 to 2014, building on previous work and proposing an updated database. Where established, the root cause for unrest is always magmatic; none was purely hydrothermal or tectonic. An interpretive classification of unrest invokes two spectra—compositional (mafic to felsic) and the state of magma conduits feeding from the magma reservoir(s) to the surface (from fully plugged, through semiplugged, to open). Magma and gas in open conduits can rise and erupt freely; magma in semiplugged conduits erupts less frequently yet still allows some gas to escape; plugged conduits allow neither magma nor gas to escape. Unrest in mafic calderas is subtler, less pronounced, and repeated, especially with open systems, ensuring the continuous, aseismic, and moderate release of magma. Plugged felsic calderas erupt infrequently, anticipated by isolated, short and seismically active unrest. Semiplugged felsic calderas also erupt infrequently and are restless over decades or centuries, with uplift, seismicity, and degassing and, on the longer-term, resurgence, suggesting repeated stalled intrusions. Finally, the expected advances in better understanding caldera unrest are discussed.

Key Points

Review of caldera unrest between 1988 and 2014
Definition of main unrest types at calderas
Original model to understand caldera unrest

1 Introduction

The most important challenge for modern volcanology is to forecast eruptions; this involves the evaluation, with some degree of confidence, of the likely hazard from an impending eruption [e.g., Sparks, 2003]. Fortunately, almost all eruptions are preceded by anomalous, usually elevated seismicity, deformation, and gas emission of the volcano—called unrest and usually lasting from a few hours to several years. Commonly reported forms of unrest are local earthquake swarms, uplift, subsidence, tilt, ground fissuring, changes in the gravity field, changes in the temperature of soil, water or gas, and changes in fumarolic activity (Figure 1) [Newhall and Dzurisin, 1988, and references therein; Sandri et al., 2004; Moran et al., 2011; Acocella, 2014]. All eruptions are thus preceded by unrest, which may show variable features from volcano to volcano and even within the same volcano. However, much unrest, even when accompanied by dramatic variations in the monitoring parameters, has not been followed by any eruption, as, for example, at Campi Flegrei (Italy) during 1982–1984, or at Long Valley (California) during 1978–2005. Since only some unrest culminates in an eruption [e.g., Newhall and Dzurisin, 1988; Sparks et al., 2012; Phillipson et al., 2013]. Unrest increases the probability of an eruption but does not push that probability to 1.0. Understanding unrest is not only important to forecast impending eruptions but also allows one to unravel the behavior and operating principles within a volcano, which is the fundamental and intermediate step to forecast eruptions.

Details are in the caption following the image — **Figure 1**
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Processes involved in caldera unrest. The emplacement of magma (sill) may induce fracturing and thus variations in the permeability of the host rock, also resulting in fluid migration and, in turn, pressure variation within a hydrothermal system. At the surface, these processes manifest themselves through seismicity, deformation, and degassing deviations from the baseline. Dike intrusions can also induce unrest across a caldera, especially if they mechanically compress and raise pore water pressures in a confined aquifer.

Calderas are probably the most complex, metastable, and dangerous type of volcano, for both humans (many large urban areas lie nearby or within calderas) and environment (including any form of life, atmosphere, climate, and infrastructures). Conversely to many volcanoes, characterized by the development of a conical edifice, at the surface calderas can be recognized by a topographic depression, from ~1 km to several tens of kilometers wide and up to a few kilometers deep, formed by the subsidence into a partly drained magma reservoir. Subsidence results from an eruption or lateral intrusion of magma [Scandone, 1990; Lipman, 1997; Cole et al., 2005; Acocella, 2007]. At depth, a caldera is often characterized by a large (reaching >10³ km³ in volume), long-lived (up to 10⁶ years), geometrically complex, heterogeneous and active magmatic system. Not every caldera can be associated with a significant subaerial eruption, as, for example, recently witnessed during the caldera collapses at Miyakejima (Japan) in 2000 or at Piton de la Fournaise (Réunion) in 2007 [Geshi et al., 2002; Michon et al., 2007]; however, the opposite is true so that all the largest eruptions on Earth, reaching up to 10³ km³ of erupted magma, are associated with the development of calderas [Scandone, 1990]. The largest caldera eruptions can affect global climate [e.g., Robock, 2000; Self, 2006]. Because of this extreme destructive potential, calderas are well studied and monitored and an important portion of the volcanological community has been trying to understand their behaviors [e.g., Newhall and Dzurisin, 1988, and references therein]. However, understanding the state of the magmatic system of calderas is commonly complicated by the presence of an active, shallower hydrothermal system(s), hosting heated and pressurized magmatic and nonmagmatic fluids (Figure 1). Variation in the state of the fluids within the hydrothermal system, even if not induced by the underlying magmatic system, may be still accompanied by seismicity, surface deformation, and degassing and thus detected by a monitoring network and indicating unrest; in addition, the hydrothermal system may undergo processes deeply modifying its permeability, including mineral dissolution and precipitation, the alteration of primary minerals, and formation of secondary hydrothermal assemblages. Nonmagmatic unrest may be exclusively related to variations in the permeability of the fracture network hosting the hydrothermal system following regional tectonic earthquakes, as, for example, suggested by some authors for Long Valley, Rabaul (Papua New Guinea), and Campi Flegrei [e.g., Hill et al., 2002; Lupi et al., 2015]. In addition, important variations within the regional stress field may directly affect the state of the magmatic system itself, inducing variations in the monitoring parameters of calderas at the surface, as observed during the 2011 Tohoku earthquake, Japan [Takada and Fukushima, 2013]. Therefore, unrest at calderas may be triggered by (a) variations within the pressure and/or volume of the magmatic system due to magma emplacement and/or ascent, (b) pressure and/or volume variations within the hydrothermal system, (c) coseismic or postseismic variations within the stress field induced by regional earthquakes, and (d) any combination of the above mentioned mechanisms. Multiple triggers for unrest, as well as any interaction between the variations in the state of the magmatic and the hydrothermal systems, constitute a further difficulty in understanding the processes behind unrest at a caldera. Indeed, different processes can produce similar symptoms of unrest: an example is the still-debated role of the hydrothermal and magmatic systems of the 1982–1984 noneruptive unrest at Campi Flegrei [e.g., Battaglia et al., 2006; Amoruso et al., 2008].

Most of our knowledge on caldera unrest has been reported and summarized in Newhall and Dzurisin [1988]. Their study showed that unrest occurs at nearly 20 calderas in a typical year, and more than half of 225 large Quaternary caldera volcanoes have experienced historical unrest, accompanied by seismicity, ground deformation, and degassing. Most seismic events in calderas are discrete earthquakes with M < 3, shallower than 15 km. Some result from brittle failure of the surrounding rock in response to magma intrusion; others reflect release of tectonic stress, shear of viscous magma along conduit walls, magma explosions, or collapse following subsurface magma flow. Ground deformation can occur on all scales. It can be sufficiently dramatic to be witnessed directly, but normally, rates are just millimeters or centimeters per year and can be detected only through leveling or gravity surveys, GPS or interferometric synthetic aperture radar (InSAR). Subtle, years-to-decades-long uplift has occurred within several large calderas, including Campi Flegrei, Rabaul, Aira and Iwo-Jima (Japan), Long Valley (California, USA), Yellowstone (Wyoming, USA), Kilauea, and Mauna Loa (Hawaii, USA). A common pattern is dome-like inflation, with maximum uplift centered within the caldera. Whereas deformation may occur throughout the caldera, most seismic activity is restricted, for example, to caldera margins. However, seismicity, uplift, increased thermal activity, and eruptions at a caldera do not always share a common center, and the centers may shift during unrest. Variations may occur in the composition or flux of fumaroles or hot springs or, more indirectly, the level and composition of groundwater or caldera-lake water. Common changes are increases in total discharge, in the discharge and proportion of acid gases (SO₂, H₂S, HCl, and HF), and in CO₂ emission. Temperature changes of caldera lakes have been reported, and the temperature of fumaroles may change by several hundred degrees. Caldera unrest can be intermittent, posing a further challenge for eruption forecasting. It may persist for weeks or centuries at large calderas, such as Aira, Campi Flegrei, Rabaul, and Iwo-Jima. The activity may wax or wane several times before culminating in an eruption, or a shallow intrusion, or returning to quiescence. Seismicity commonly occurs in repeated swarms, and uplift may alternate with subsidence. For example, subsidence at Yellowstone caldera during 1985–1987 followed decades of net uplift. Ground deformation may be episodic, as indicated by stepped terraces at Toba (Sumatra, Indonesia), Campi Flegrei, and Iwo-Jima calderas. Unrest a few days or hours before an eruption is typically characterized by dramatically increased uplift rates to meters per day, as well as the increase of the number of earthquakes, becoming progressively shallower. These features are commonly attributed to the propagation of the feeder dike from a magma reservoir to the surface.

The compilation and synthesis of Newhall and Dzurisin [1988] was comprehensive for its time, but many of the data in that compilation predate modern monitoring and are thus descriptive and qualitative, with very limited quantitative information. The boosting of monitoring techniques at many volcanoes worldwide in the last decades permitted the collection of much more quantitative and thus precise data on caldera unrest. These data on unrest at calderas in the last decades are dispersed and fragmented through publications, reports, and websites. The collection and systematic and critical review and analysis of such information may provide crucial insights for identifying general types, establishing patterns, thresholds, and relationships and, ultimately, understanding unrest processes.

The authors of this study have collected data on recent caldera unrest from 1988 to 2014. The aim of this study is twofold: (1) to update details of recent caldera unrest and (2) offer a modern framework to better understand caldera unrest and forecast eruptions at calderas. Our work should be further improved by a modern probabilistic approach and more detailed studies devoted at understanding the processes behind the observables.

2 The Database

There are 446 known calderas worldwide [Geyer and Martí, 2008; Sobradelo et al., 2010], including ~225 active in the Quaternary [Newhall and Dzurisin, 1988]. Of these, 97 have been active during the Holocene [Geyer and Martí, 2008; Siebert et al., 2010]. The pre-1988 observations and monitoring information on caldera unrest have been described in Newhall and Dzurisin [1988] and therefore are not included here. Monitoring data from 1988 to 2014 are available on 42 of the 97 active calderas. The necessity of a continuous financial support (for both equipment and personnel), as well as any inaccessibility or remoteness of a caldera, may hinder a regular monitoring; therefore, these 42 calderas have been mainly monitored because of their continuous activity, proximity to inhabited areas, potential hazard, or availability of funds. In addition to these 42 calderas, there are 8 calderas which have been sporadically monitored with synthetic aperture radar (SAR) data over ~6 years, some of them showing unrest [Biggs et al., 2009, 2011], but owing to poor and fragmented information, these latter calderas are not considered here. The monitoring information on unrest at the 42 calderas active from 1988 to 2014 is derived from different sources, including publications (~450 sources), observatories reports (~200), and Smithsonian Institution Global Volcanism Program reports (~150). Caldera names are here reported as in Siebert et al. [2010]; calderas not included in that reference are referred to by their most commonly used name. We considered calderas in any tectonic setting: convergent plate boundaries, divergent plate boundaries, and intraplate. Convergent plate boundaries include calderas in continental collision, island arc collision, oceanic back-arc rifting, and continental back-arc rifting. Divergent plate boundaries include calderas along oceanic ridges and continental rifts. Intraplate settings include hot spot calderas (Figure 2).

The monitoring information considered here documents variations in the geophysical, geochemical, and geodetic parameters. Some we term “indicators,” defined as any monitoring parameters whose value and/or rate variation is significant for interpreting unrest. Geophysical indicators mostly include seismicity and microgravity changes. Seismicity, recorded in 95% of 42 examined cases, is the most frequent indicator; it includes regional earthquakes, volcano-tectonic earthquakes, long-period events, and tremor. Microgravity changes are a much less common indicator, being measured only in a few cases (12%). Geodetic indicators, commonly but not exclusively highlighting inflation, have been recorded in 80% of the 42 examined cases. These indicators are usually obtained by GPS, InSAR, and leveling data. Geochemical/hydrothermal indicators include degassing and temperature variations from fumaroles and physical-chemical changes at the water of crater lakes. In particular, degassing variations, observed in 71% of the 42 examined cases, involve gas plumes from vents, changes in the fumarolic activity, and thermal anomalies. The less common physicochemical changes at crater lakes, observed in 10% of cases, include variations in water level, pH, temperature, and increase in gas discharge.

Here we define unrest as an anomalous state of a caldera, when at least one of the aforementioned indicators deviates from the baseline. Moreover, we define a “restless phase” as a general and prolonged period of unrest, consisting of continuous deviations from the baseline of any of the indicators, and lasting from several years to centuries; the identification of century-long unrest depends on the availability of longer-term measurements or observations on increased seismicity, anomalous degassing, or variations in the Earth's surface. Within the restless phase, any shorter (hours to a few years) period of intensification of the rates or values of any indicator is referred to as an “unrest episode.” Therefore, a restless phase may consist of several unrest episodes. For example, Rabaul caldera has been restless from the early 1970s, with at least three major unrest episodes, characterized by a marked intensification of seismicity and uplift, in the 1980s and 1990s and again in 2006.

For each unrest episode at a given caldera, we first considered the caldera descriptive parameters (Figure 3). These include the caldera name, universal transverse Mercator coordinates, caldera maximum and minimum diameters, magma composition, preeruptive unrest duration (eruptive unrest only), unrest duration, time from previous unrest, and time from previous eruption. We then focused on the unrest indicators (Figure 3). The geophysical data include the location (with respect to the caldera center) and width of the area undergoing seismicity, the maximum number of seismic events, the maximum magnitude of the earthquakes, the identification of volcano-tectonic and long period events, the frequency of tremors, and microgravity data. The geodetic data include the location of the deformation (with respect to the caldera center), the detection of uplift or subsidence, the rate of deformation, the total accumulated deformation, and any tilt measurement. The geochemical data include the position (with respect to the caldera center) of the measured anomalies, changes in the maximum temperature, variations in the chemical composition (mainly CO₂ and SO₂), pH changes, and any geysering event.

In an electronic supplement to this paper we consider and summarize these monitoring data from the 42 calderas, providing a qualitative analysis and trying to detect first-order, common and representative types. These 42 calderas are subdivided into those without eruptions (11 calderas) and those where unrest culminated in eruptions (31 calderas). The total number of unrest episodes for both the nonerupting and the erupting calderas is 166; of these, 110 culminated in an eruption. This implies that a caldera may have undergone several episodes of unrest, some of which were noneruptive. We distinguish mafic calderas as those characterized by a predominant basaltic or primitive composition, versus felsic calderas, with more evolved compositions including trachytes, andesites and rhyolites; composition affects magma density, viscosity, and the accumulation of gas. We distinguish conduits in which magma solidifies between eruptions (plugged), remains molten and fluid between eruptions (open), or reaches an intermediate state (semiplugged; viscous and often unable to erupt but still allowing gas to pass). Open conduits allow magma and gas to ascend freely; semiplugged conduits allow some gas to ascend and occasional eruptions; and plugged conduits trap both gas and magma. Calderas with open conduits exhibit less intense seismic precursors to eruptions than those with plugged conduits.

The general relationships between the calderas (28 mafic and 14 felsic) and the unrest episodes (noneruptive or eruptive) are summarized in Table 1. Phreatic eruptions have been included in the eruptive unrest category, assigning the phreatic eruption a volcanic explosivity index (VEI) = 1.

Table 1. Relationships Between Noneruptive and Eruptive Unrest and Mafic and Felsic Calderas

The collected information on the unrest indicators may be significantly inhomogeneous and discontinuous as a function of the type, modality, and duration of the acquisition of the monitoring data. Furthermore, the monitoring of the geochemical data may be uneven in terms of measured parameters (e.g., SO₂, CO₂, H₂O/CO₂, pH, and temperature of crater lakes) and unit of measurement. Lack of information on one or more indicators may result from a real absence of any geophysical, geodetic, or geochemical deviation from the baseline or, alternatively, from the effective lack of any monitoring. These different conditions, when known, are referred to as CQ 7(confirmed quiescence) or UQ (uncertain quiescence) in our database, which is included in the supporting information.

Of the 42 calderas in this study 12% were restless continuously from 1988 to 2014 (our cutoff date); 69% alternated unrest with confirmed quiescence periods (CQ); and 19% alternated unrest with uncertain quiescence periods (UQ), as these calderas have not been adequately monitored.

All monitored calderas have shown at least one episode of unrest between 1988 and 2014.

3 Types of Caldera Unrest

Below we summarize the post-1987 unrest activity of the most representative of the 42 monitored calderas. These include the best studied calderas, usually with the most complete data set, covering the variable compositional and evolutionary spectrum, from mafic to felsic calderas and from small-sized calderas on stratovolcanoes to large and long-lived magmatic systems.

3.1 Unrest Without Eruptions

3.1.1 Mafic Calderas: Representative Types

3.1.1.1 Askja and Krafla

Askja lies along the oceanic ridge of Iceland; it has been active since 200–300 ka and includes three main calderas. The largest, properly named Askja caldera, formed in early Holocene and is 8 km wide; an older and less pronounced caldera, almost completely filled with lava, is located along the northern edge of Askja caldera (Figure 4). The third, most recent caldera formed during the rifting event of 1874–1876 and is filled by the Öskjuvatn Lake, associated with the late March 1875 explosive eruption of mixed basalt-rhyolite magma. Öskjuvatn caldera is E-W elongated, 6 × 5 km wide, 230 m deep and with widespread hydrothermal activity. After continued significant subsidence, the caldera took >40 years to reach a stable state and since 1932 the collapse volume has remained constant within error. The most recent eruption, in 1961, opened an E-W trending fissure along the rim of Askja caldera [Sigmundsson, 2006, and references therein; e.g., Hartley and Thordarson, 2012].

From 1973 to 1998, a cluster of earthquakes, reaching M 4.2, occurred beneath the SE part of the volcano and along the NW rim of the main caldera; here a new phase of subsidence between 1983 and 1998 totaled >75 cm, with a rate of 50 mm/yr [Sturkell and Sigmundsson, 2000; de Zeeuw-van Dalfsen et al., 2012]. A similar continuous subsidence was detected also in the noneruptive 2000–2009 period, even though the rate lowered to 25–30 mm/yr. The subsidence was accompanied, between 1988 and 2007, by a net gravity decrease without any eruption, large earthquake, or dike injection. The subsidence is probably caused by a combination of a cooling and contracting magma chamber, at ~3.5 km depth [De Zeeuw-van Dalfsen et al., 2005, 2012, 2013; Pagli et al., 2006]. Lower crustal earthquakes (depth >12 km) occurred from 2005, whereas from 2006 seismicity moved to the upper lower crustal boundary. Between 2007 and 2009, a gravity increase occurred at the caldera center, suggesting that magma may have flowed into the shallow magma chamber at ~3 km depth, despite continued subsidence [Rymer et al., 2010; de Zeeuw-van Dalfsen et al., 2013].

In summary, in the last decades Askja has been characterized by continuous subsidence, with overall decreasing rates. The subsidence was partly accompanied by seismicity, but the depth of the source of the subsidence appears different and shallower. Askja constitutes an example of continuously deflating caldera, even more than a century after a major eruption. The deflation phase may have been ended now and possibly replaced by a new phase of magma injection in the shallow chamber.

Similar behavior to Askja has been observed in the nearby Krafla caldera. The 8 × 10 km wide Krafla caldera may have formed at ~0.1 Ma [Sigmundsson, 2006, and references therein]. A rifting episode in 1975–1984 activated most of the Krafla magmatic system. Magma flow into a shallow magma chamber caused inflation of the caldera region, interrupted by sudden deflation events due to the lateral propagation of dikes along the magmatic system. Twenty inflation/deflation cycles were recorded over the course of the rifting episode, and nine resulted in eruptions [Bjornsson et al., 1977; Sigurdsson, 1980; Buck et al., 2006, and references therein]. During the 1989–2005 postrifting period, Krafla totaled a subsidence of ~30 cm, largely due to a pressure decrease within the magma chamber; the rate passed from ~5 cm/yr in the 1989–1992 period to <0.3 cm/yr in the 2000–2005 period. The most recent subsidence has been largely attributed to fluid extraction for geothermal exploitation [Sturkell et al., 2008]. The subsidence was accompanied by minor dispersed seismicity, with M < 2, as well as net gravity decrease between 1990 and 1996, interpreted as due to drainage from a shallow magma chamber. Similar gravity decreases have not been observed during the posteruptive deflation at Askja [Rymer et al., 1998b; de Zeeuw-van Dalfsen, 2004; Sturkell et al., 2008, and references therein].

3.1.1.2 Mauna Loa

Mauna Loa, Hawaii, has been producing tholeiitic lavas for hundreds of thousands of years during its shield-building stage. The volcano has two well-developed rift zones, the NE and the SW Rift. Mauna Loa's summit caldera, Mokuaweoweo, formed ~1200 years ago as magma drained away to feed a large flank eruption 35 km away and 2500 m lower on the volcano's NE rift [Lockwood and Lipman, 1987; Barnard, 1995]. The caldera coalesced with the two main pit craters (north and south pits) and several smaller, discretely formed ones. The last eruptions at Mauna Loa occurred in 1975 and 1984. From 1990, earthquake counts presented fluctuating trends with an apparently increase since late 1990. Deformation measurements from 1991 indicate gradual reinflation of Mauna Loa's summit. This trend reversed in 1993–1994, when distances across the caldera shortened by as much as 7 cm, and leveling surveys in 1996 and 2000 measured more than 7 cm of subsidence at SE of the caldera [GVP report, 09/2002 (BGVN 27:09)]. The pattern of deformation at Mokuaweoweo reversed from deflation to inflation from May 2002, when GPS data indicated lengthening at a rate of 5–6 cm/yr and shallow earthquakes reached M 2.5 at ~3 km beneath the SW Rift. Inflation at rates as high as ~8 cm/yr on the eastern and western sides of the caldera continued until 2005, accompanied by an increase in subcrustal seismicity in 2004. The 2002 to 2005 inflation has been explained by the emplacement of a dike below the caldera and the SW Rift, at a depth of 3 to 8 km; the inferred rate of magma accumulation of 21 × 10⁶ m³/yr is almost 3 times the long-term growth rate averaged over the past 4000 years. Dike emplacement has been continuing, even though at a lower rate, in 2007 and in 2014 [Amelung et al., 2007; GVP report, 05/2012 (BGVN 37:05)]. During 2004–2010, several small long-period (LP) earthquakes and variations in gas emissions occurred to the south of the caldera.

Summarizing, Mauna Loa in the last decades was characterized by at least one major episode of unrest between 2002 and 2005, in which the emplacement of a dike was accompanied by surface deformation, seismicity, and minor degassing.

3.1.1.3 Taal

Taal is an active composite volcano in the Philippines, which includes a 16 × 27 km prehistoric caldera, partly filled by Lake Taal. The caldera results from at least four major ignimbrite eruptions between 500 and 100 ka [Listanco, 1994; Torres et al., 1995; Lowry et al., 2001]. The erupted products have a predominantly basaltic to basaltic andesite composition, though dacitic products are also erupted. As the recent activity is characterized by 50–59% of SiO₂, we include Taal in the mafic calderas section. The ~5 km wide Volcano Island occupies the center of Lake Taal, and eruptions of recent decades were concentrated on the SW flank (Figure 5) [Richon, 2003]. Seismicity, deformation, geysering, and changes in water chemistry and temperature characterize the activity of the caldera. In early 1991 high-frequency earthquakes occurred under the east side of Volcano Island, at depth of 2–5 km. In addition, the Main Crater Lake revealed an increase of acidity and a possible slight increase in water temperature. However, no inflation was observed [GVP report, 03/1991 (BGVN 16:03)]. Two other episodes of unrest, in 1992 (February–March) and 1994 (March), consisted of intense seismicity, sudden ground deformation, and changes in lake temperature and chemistry [Bartel et al., 2003, and references therein]. In February 1992, up to 21 cm of uplift in <1 day was measured, while the seismicity reached a peak of 385 high-frequency earthquakes [Bartel et al., 2003; GVP report, 02/1992 (BGVN 17:02)]. In mid-March 1994, both the north and SE sides of Volcano Island inflated 10–20 cm, with an increase of temperature of the Main Crater Lake and seismic events [Bartel et al., 2003; GVP report, 02/1994 (BGVN 19:02)]. Between 1998 and 2000 four episodes of deformation, seismicity, and geyser activity have occurred. Deflation and a geysering episode on Volcano Island (August 1998) from June to December 1998 were followed, in January–March 1999, by an inflation of ~20 mm and another geysering episode, without notable variations in seismicity. Between March 1999 and February 2000, a new episode of deflation consisted of daily geysering events, with two periods of increased seismicity [Lowry et al., 2001; Bartel et al., 2003; Zlotnicki et al., 2009]. Finally, between February and November 2000, waning geysering activity was accompanied by a slight increase in seismicity and maximum uplift of ~120 mm of the center of Volcano Island relative to the northern caldera rim. The source of 1999 deflation and inflation in 2000 were modeled as contractional and dilatational Mogi point sources centered at 4.2 and 5.2 km depth, respectively, beneath Volcano Island. The two periods of inflation observed at Taal result from episodic intrusions of magma into a shallow reservoir centered beneath Volcano Island. Subsequent deflation may result from exsolution of magmatic fluids and/or gases into an overlying, unconfined hydrothermal system [Bartel et al., 2003].

More recent unrest episodes occurred at Taal between 2008 and 2015. In late August 2008 very shallow earthquakes (<1 km depth) occurred along the north rim of the Main Crater. These were followed, in 2010, by an increase in the number and magnitude of the earthquakes, both high and low frequencies, by steam emissions from the main crater and a few millimeters of inflation. Ground inflation continued through May 2014 to March 2015, sporadically associated with seismicity, even though CO₂ emissions decreased from November 2014 to January 2015 [GVP report, 02/2011 (BGVN 36:01)].

In synthesis, unrest of Taal caldera in the last decades has been episodic, with an overall good correlation between the rates of the observed indicators. However, in the 1999–2000 period, the relationship between deformation and other indicators was ambiguous (Figure 5), because in 1999 seismicity occurred during the yearlong deflationary trend and lower seismicity, while the lack of geysering events in 2000 was accompanied by inflation.

3.1.2 Felsic Calderas: Representative Behaviors

3.1.2.1 Campi Flegrei

Campi Flegrei is the main active Neapolitan volcano (Italy), between Somma-Vesuvio and Ischia. Campi Flegrei has been active in the last 50 ka at least, and its main feature is a ~15 km wide caldera structure formed during two major eruptions: the Campanian Ignimbrite eruption, at 39 ka, and the Neapolitan Yellow Tuff (NYT) eruption, at 15 ka (Figure 6) [Scandone et al., 1991; Orsi et al., 1996; Deino et al., 2004; Giacomelli and Scandone, 2012]. The former eruption created a pair of nested ring faults, reactivated by the latter eruption [Acocella, 2008]. Intracaldera activity occurred episodically along the border of the caldera, after the two major events [Rosi et al., 1983; Rosi and Sbrana, 1987]. Di Vito et al. [1999] suggest that there were flare-ups of activity, called epochs, alternating to periods of quiescence. Epoch I lasted from 12 to 9.5 ka, giving rise to 34 explosive eruptions. During Epoch II, dated between 8.6 and 8.2 ka, six explosive eruptions took place. Epoch III lasted from 4.8 to 3.8 ka and produced 16 explosive and 4 effusive eruptions. A significant uplift of the central part of the caldera occurred after the collapse following the NYT eruption. An en echelon pattern of normal faults in the northern sector of the Gulf of Pozzuoli [Colantoni et al., 1972] borders an uplifted portion of the caldera floor which, on land, is marked by a raised marine terrace, called “La Starza,” presently at 40 m above sea level (asl) [Cinque et al., 1985]. A higher coastal cliff at a maximum elevation of 60–65 m borders the northern side of the La Starza terrace [Cinque et al., 1997]. The beginning of the uplift phase is not known with certainty; the oldest marine sediments at the base of the terrace have been dated at ~10 ka [Cinque et al., 1997]. The fossils of the upper marine deposits suggest that the maximum uplift (considering the depth of deposition) may be of ~65 m [Ciampo, 2004]. The last eruption of Monte Nuovo occurred in 1538, after a repose period of circa 2500 years with seismicity and uplift progressively increasing in the 100 years before [Orsi et al., 1996; Di Vito et al., 1999; Guidoboni and Ciuccarelli, 2011; Giacomelli and Scandone, 2012].

The most recent evolution has been characterized by short-term ground deformation, culminating in at least three major unrest episodes between 1950 and 1952, 1969 and 1972, and 1982 and 1984, with uplift of ~0.7, ~1.7, and ~1.8 m, respectively [Del Gaudio et al., 2010]. Of these, the best monitored 1982–1984 unrest was accompanied by M < 4.2 seismicity in the central part of the caldera, clustered at depth <3–4 km below the central Pozzuoli area, and by newly formed surface fractures. The 1982–1984 unrest may have resulted from the pressurization of a hydrothermal system, probably under the input of new magma or magmatic gases [Battaglia et al., 2006; Chiodini et al., 2003; Bodnar et al., 2007; Amoruso et al., 2008]. After the 1982–1984 unrest, the caldera underwent marked subsidence, reversing almost 1 m of uplift by 2004. The subsidence was spaced out each few years by moderate, short and localized uplifts (<5 cm, centered in the Solfatara crater area next to Pozzuoli), accompanied by seismicity and increased hydrothermal activity. However, after some further variation in the composition and flux of the fumarole gases, measured since 2000 [Chiodini et al., 2010], from 2005 the central part of the caldera also underwent a progressive uplift, totaling nearly 28 cm to early 2015 (Figure 6) [De Martino et al., 2014]. This uplift has been partly accompanied by shallow (<4 km) microseismicity and a marked increase in the pressure and temperature of the hydrothermal system and in the flux of the magmatic gases (Osservatorio Vesuviano data, http://www.ov.ingv.it/ov).

The onset of this change dates back to 2000–2005, when compositional variations in the gases and relatively deep long-period seismicity suggested ascent of magmatic fluids toward the shallower hydrothermal system [Chiodini et al., 2012a; Petrillo et al., 2013]. The entire post-1980 deformation at Campi Flegrei results from the activity of a deeper magmatic source, at 4 km depth below Pozzuoli, and a shallower hydrothermal source, at 2 km below the Solfatara area [Amoruso et al., 2014]. A continuous transfer of magmatic fluids from the deeper to the shallower source has been responsible for the observed dynamics in the last decades [Chiodini et al., 2015].

In synthesis, repeated unrest episodes within Campi Flegrei are of a type induced by the activity of and interaction between a deeper magmatic and a shallower hydrothermal source. The unrest episodes are often characterized by an overall coupling in the increase of surface deformation, degassing, and seismicity, though the last unrest was accompanied by poor seismicity. The ongoing unrest was at first geochemically detected and then followed by inflation [Newhall and Dzurisin, 1988; Chiodini et al., 2012a].

3.1.2.2 Long Valley

Long Valley Caldera (eastern California) is the largest structure in the Long Valley Caldera-Mono Craters volcanic field that includes Mammoth Mountain and the Mono-Inyo volcanic chain [Hill et al., 2003; Hill, 2006]. The caldera is an E-W elongated elliptical depression ~32 × 17 km (Figure 7) [Bailey et al., 1976; Bailey, 1989]. An uplift of ~500 m in the last 600 ka in the west central portion of the caldera produced a resurgent dome dissected by northwest trending faults [Bailey et al., 1976; Bergfeld et al., 2006]. Postcaldera volcanism progressively decreased until Holocene, developing the Inyo Chain Sequence, culminating in the Mono-Inyo eruption of 1350 A.D., to the NW of the caldera [Romero et al., 1993; Hildreth, 2004; Mahood et al., 2009; Seccia et al., 2011].

From 1975 to 1999 the resurgent dome uplifted 75 cm. This uplift has been, in general, related to the inflation of a magmatic source at a mean depth of 5.9 km beneath the dome [Tizzani et al., 2009]. In detail, the inflation of the resurgent dome between 1989 and 1991 was associated with local seismicity; the onset of deformation preceded the onset of increased earthquake activity by more than 2 months. On the SW caldera rim (Mammoth Mountain area), diffuse CO₂ degassing in 1990 was associated with local seismicity [Langbein et al., 1993]. Two longest swarms occurred during 24–27 March 1991 and late November to early December 1993. The March 1991 swarm included >1000 detected earthquakes, with 22 M > 3 earthquakes [Hill, 2006]. Immediately after the 28 June 1992 Landers earthquake, 400 km south of Long Valley, seismicity and deformation increased significantly, suggesting that these could have reflected movement of bubbles within the magma body [Linde et al., 1994]. Strong surface waves from the Landers earthquake also triggered a transient, caldera-wide uplift that reached a peak uplift of ~5 mm 5–6 days after the Landers earthquake. The June 1992 swarm itself included >250 located earthquakes [Hill, 2006]. In contrast to the general tendency of caldera earthquake swarm activity to follow the uplift rate of the resurgent dome, one of the strongest earthquake swarms in the caldera occurred in March and April of 1996, as inflation of the resurgent dome was slowing [Hill et al., 2003]. During 1997–1998, the deformation rate was an order of magnitude higher than the previous 3 year average, increasing exponentially and then decreasing exponentially with a similar rate [Langbein et al., 1993; Langbein et al., 1995; Marshall et al., 1997; Newman et al., 2001; Langbein, 2003]. Additional rapid episodic uplift occurred between 2002 and 2003, associated with low background seismicity and related to a source at depth of 7.5–13.5 km [Feng and Newman, 2009]. The maximum rate of uplift was similar to that observed during the uplift episodes in 1989–1990 and 1997–1998; in fact, all three episodes began immediately after a short period of seismic quiescence, with background seismicity falling to levels below background levels following the prior uplift event. After 2003, slow subsidence occurred between 2004 and 2007, followed by slow uplift between 2007 and 2009. The source of the latter, producing a low volume change, was located at depths of 6–8 km and can be related to a mixture of hydrothermal and partial-melt magma [Feng and Newman, 2009; Liu et al., 2011].

In synthesis, during the last decades the resurgent dome area in the Long Valley caldera has experienced episodic uplift, with variable and, in general, decreasing rates. With minor exceptions, as in 1996, the uplift has been coupled with an increase in the seismicity, both within the caldera and the nearby Sierra Nevada block [Hill, 2006].

3.1.2.3 Yellowstone

The Yellowstone Plateau covers an area of ~6500 km² and was formed during three supereruptions at ~2.05, ~1.3, and ~0.64 Ma [Christiansen, 2001]. The last eruption produced the current 40 × 60 km wide Yellowstone caldera, one of the most intense regions of magmatic, seismic, and hydrothermal activity on Earth [Dzurisin and Yamashita, 1987; Dzurisin et al., 1990, 1994; Christiansen, 2001; Wicks et al., 2006; Vasco et al., 2007; Aly et al., 2009; Hurwitz and Lowenstern, 2014]. Since the last supereruption and after the uplift that formed the two Sour Creek (SC) and Mallard Lake resurgent domes, 50 basaltic and rhyolitic nonexplosive eruptions of lava or less violent explosive eruptions have occurred, the most recent at ~70 ka (Figure 8) [Christiansen, 2001]. The current activity of Yellowstone is more or less continuous subsidence, uplift, and thermal activity, with occasional episodes of higher seismicity [Aly and Cochran, 2011; Hurwitz and Lowenstern, 2014].

Between 1923 and 2004 several episodes of uplift and subsidence have occurred, with the highest average rates of 1–2 cm/yr focused on the resurgent domes. In particular, between 1986 and 1995, subsidence in the caldera center, with a maximum of ~8 cm and a velocity between 1.6 and 2.7 cm/yr, was coeval to an uplift detected near the Norris Geyser Basin (NGB), with an average rate of 0.5–0.8 cm/yr [Chang et al., 2007; Aly and Cochran, 2011]. In 1996–2000 there was an uplift of ~6.8 cm in the NW edge of the caldera near NGB with a rate of ~1.5 cm/yr, and minor uplift within the caldera. Between 2000 and 2004 a subsidence of the caldera of 2.8 cm, with a rate of ~0.9 cm/yr, corresponded to an uplift of the NGB of ~2.6 cm [Chang et al., 2007; Aly and Cochran, 2011]. From the second half of 2004 to 2006, inflation progressively increased toward the caldera center, along its major NE-SW axis, with the highest rate (~7 cm/yr) at SC; contemporaneously, NGB subsided 3.7 cm/yr. During this period, seismic activity focused near the northern border of the caldera, between the inflating and deflating zones. This deflation of the NGB was probably triggered by the redistribution of hydrothermal fluids as a result of caldera inflation [Chang et al., 2007]. During 2008–2009, the rate of uplift across the Yellowstone caldera slowed to ~3 cm/yr and subsidence virtually ceased at the NGB [Aly and Cochran, 2011].

In synthesis, Yellowstone underwent three major episodes of unrest in the last decades: 1986–1995, 2000–2004, and 2004–2009. The 1988–1995 and 2000–2004 periods were characterized by subsidence within the caldera, not always associated with high seismicity, and uplift in the NGB. Conversely, the unrest episodes between 2004 and 2009 showed uplift in the caldera and subsidence on the NGB. Overall, Yellowstone marks the restless activity of a complex magmatic system, where the deformation of the main hydrothermal area outside the caldera (NGB) appears generally anticorrelated to the caldera center, with fluid transfer from one area to the other. Even if the most seismic activity seems correlated to the deformation in the NGB area, there was overall coupling between seismicity in the NGB area and uplift of the Sour Creek and Mallard Lake resurgent domes [Chang et al., 2007, 2010].

3.1.2.4 Colli Albani

Colli Albani Volcano (Italy) has erupted mainly leucititic magmas between 0.6 and 0.02 Ma and is regarded as one of the most explosive mafic volcanoes, even though its morphology and predominant eruptive style match those of felsic calderas; for these reasons, the volcano is included in the felsic calderas section. The most distinctive feature of Colli Albani is an 8 km wide caldera, formed in repeated eruptions between 600 and 355 ka (Figure 9) [Funiciello and Giordano, 2010, and references therein]. Phreatomagmatic activity formed several maars and tuff cones on the western and northern slopes between 200 and 23 ka. Holocene phreatic activity in the Albano maar occurred 5–6 ka B.P. and probably in the fourth century before the Common Era. High lake levels and catastrophic withdrawal of the Albano maar lake since prehistoric age are possible indicators of sudden variation of CO₂ flow and upwelling of hydrothermal fluids [Funiciello et al., 2003; Carapezza et al., 2008]. Seismic tomography identifies a low-velocity region, perhaps still hot or partially molten, >6 km beneath the youngest craters, and a high-velocity region, probably a solidified magma body, beneath the older central volcanic construct [Chiarabba et al., 1997].

Colli Albani has recently experienced moderate-intensity earthquakes, seismic swarms, gas emissions, and ongoing uplift. An uplift of ~30 cm (average of 7 mm/yr) was detected by leveling surveys performed in the period 1951–1994 along a leveling line that crosses the more recently active western flank, where seismic activity focuses at depths of 3–6 km [Amato and Chiarabba, 1995]. More recent space-based GPS and InSAR data confirm that this uplift persists also after 1994, with rates between 3 and 6 mm/yr, and is distributed in a wide area around the youngest craters of Albano and Nemi. The measured deformation can be linked to one or more magmatic sources, 4.6 to 7 km deep, below the western flank [Salvi et al., 2004; Riguzzi et al., 2009; Anzidei et al., 2010]. Here there has been widespread gas emission (mainly CO₂ of magmatic origin and subordinately H₂S) during Holocene and continuing to the present day [Tuccimei et al., 2006; Carapezza et al., 2008, 2010, 2012]. In particular, Lake Albano, on the western flank, was affected by a large CO₂ input that coincided with the last important seismic swarm in 1989, suggesting an intimate relationship between the addition of deep-originated CO₂ to the lake and seismic activity. The total mass of dissolved CO₂ decreased from ~5.8 × 10⁷ kg in 1989 to ~0.5 × 10⁷ kg in 2010, following an exponential decreasing trend [Chiodini et al., 2012b].

Seismicity at depths between 5 and 3 km below the western flank, as observed in 1989–1990, may be related to the inflation of a magmatic source, responsible for an increase of Coulomb stress >0.5 MPa in the overlying crust: >85% of fault plane solutions are consistent with the stress perturbations induced by this magmatic source, suggesting that microearthquakes are promoted by elastic stress changes [Feuillet et al., 2004].

In synthesis, Colli Albani is a caldera undergoing long-term unrest with an overall coupling between geodetic, seismological, and geochemical indicators. However, the observed rates of the monitoring indicators at Colli Albani do not reach the high values of the similarly restless Campi Flegrei, Long Valley, and Yellowstone calderas.

3.1.2.5 Santorini and Nisyros

The South Aegean volcanic arc results from the subduction of the East Mediterranean oceanic lithosphere below the European continental plate [e.g., Papanikolaou, 1993; Ersoy and Palmer, 2013, and references therein]. Recent volcanoes with calderas formed along this arc are Santorini, Nisyros, and Kos.

Santorini (Thera, Greece), in the center of the South Aegean volcanic arc, is the most active volcano in the Aegean Sea and consists of five islands, both in the center (Kameni Islands) of the 6 to 10 km wide caldera and along its sides (Figure 10) [Stiros et al., 2010]. Volcanism in the area started at 2 Ma, the most important event probably being the caldera-forming Minoan eruption at ~3.6 ka, for which the recharge of the magma reservoir occurred during the century before, while mixing between different magma batches was still taking place during the final months [Druitt et al., 2012]. The Kameni Islands, a dome complex in the caldera center, were formed during the last 2000 years along NE-SW trending regional normal fault systems composing a rift oblique to the Aegean arc [Heiken and McCoy, 1984; Newhall and Dzurisin, 1988].

Since its last eruption in 1950, Santorini remained in an overall dormant state. In the period 1992–2010 a general and gradual deflation of 5–6 mm/yr occurred over Nea Kameni, in the caldera center, without seismic activity [Foumelis et al., 2013]. However, on the northern part of the caldera (between Nea Kameni and Therasia) up to 10 cm of gradual and aseismic inflation has been detected between June 1994 and May 2003; a Mogi source was inferred to be active between the Nea Kameni and the Therasia islets along a major tectonic structure [Stiros et al., 2010]. Between January 2011 and January 2012 the entire caldera showed signs of unrest, with increased microseismic activity, significant ground uplift, and variations in the composition of fumarolic gases. Earthquakes with magnitude between 1.0 < M_L < 3.2 focused within the caldera at 1–6 km depth, probably under the NW-SE extensional stress field that controls the distribution of seismic activity promoted by the unrest [Newman et al., 2012; Feuillet, 2013; Lagios et al., 2013]. The nonlinear uplift focused in the northern part of the caldera (50–120 mm/yr) with increasing rate in Nea Kameni [Lagios et al., 2013]. The H₂ and CO₂ concentrations in the central part of the caldera increased from ~0.01 to ~98.7 mmol/mol and from ~400 to ~500 mmol/mol, respectively; this increase is explained by a convective heat pulse from depth associated with the seismic activation of the NE-SW oriented Kameni fault, possibly triggered by either injection of new magma or increased permeability of the volcanic plumbing system [Tassi et al., 2013]. Joint inversions of SAR and GPS velocities using spherical and spheroidal magmatic source types indicate their location offshore at ~1 km north of Nea Kameni and between 3.5 and 6 km depth [Papageorgiou et al., 2012; Newman et al., 2012; Parks et al., 2012; Foumelis et al., 2013; Lagios et al., 2013]. However, detailed studies on the distribution of the microseismicity during the unrest also might be interpreted as concurrent magma pulses distributed along offset, regional tectonic structures [Saltogianni et al., 2014].

In synthesis, the recent unrest at Santorini, with a coupling between seismic, geodetic, and geochemical monitoring indicators, has been due to the rapid injection of magma into the shallow system after long repose times. Rapid assembly before eruption of the liquid magma bodies that fuel giant caldera-forming eruptions has been here documented [Druitt et al., 2012], but the new observations suggest that the same process may be responsible for smaller eruptions too [Hooper, 2012].

A similar isolated episode of unrest, after several decades of quiescence occurred on nearby Nisyros Island (Greece). This unrest, triggered by a magmatic injection and amplified in its manifestation by the hydrothermal system, lasted longer than the one on Santorini but was similarly characterized by an overall coupling in the increase in the rate of all the monitoring indicators. The Quaternary Nisyros volcano consists of an 8 km wide truncated cone, formed between 150 and 25 ka, and a 4 km wide summit caldera related to a plinian eruption <20 ka. Postcaldera activity has been characterized by dacitic lava domes inside and outside the Lakki Plain, representing the central eastern caldera floor [Limburg and Varekamp, 1991; Bachmann et al., 2011]. The western sector of the Lakki Plain was affected by at least 13 historical hydrothermal eruptions, the most recent occurring in 1871–1873 and 1888, accompanied by violent earthquakes, gas detonations, steam blasts, and mudflows [Marini et al., 1993]. Faults cluster in left-lateral NE-SW and extensional N-S systems. The NE-SW preferred strike of the Lakki faults and of the mineral-filled veins, as well as the distribution and NE-SW elongation of the hydrothermal craters, indicates that tectonics plays a major role in controlling the fluid pathway. The same NE-SW trend is depicted by CO₂ anomalies, indicating a structural control on hydrothermal degassing [Caliro et al., 2004].

The 1995–1998 unrest is the first instrumentally documented at Nisyros. Differential InSAR (DInSAR) analysis between May 1995 and September 2000 suggests that deformation has occurred at least since 1995. Shallow seismic activity between 1995 and 1998 was highly clustered, both in space and time. From November 1995, 113 M_s ≥ 4.1 earthquakes occurred, with >50% at ≤10 km of depth, the largest one occurring on August 1997 with M_s 5.3 [Papadopoulos et al., 1998; Lagios et al., 2005]. A GPS network deployed between 1997 and 2002 measured an uplift ranging from 14 to 140 mm and horizontal displacements from 13 to 53 mm. A two-source Mogi model combined with assumed motion along the N-S trending Mandraki Fault fits the observed GPS and DInSAR deformation. This assumes sources at a depth of 5.5 km NW of the center of the island and at 6.5 km offshore, ESE of the nearby Yali Island, to the NW of Nisyros [Lagios et al., 2005]. Between 1999 and 2000, an aseismic deflation of 15 mm was recorded. Between 2000 and 2002, uplift and slight seismicity restarted, accompanied by visible ground deformation, expressed in the reactivation of the Mandraki Fault and a 600 m long N-S trending irregular rupture in the caldera floor [Lagios et al., 2005]. In this period, the geochemical data also showed a dramatic increase in H₂S/CO₂ and a decrease in CH₄/CO₂ for all fumaroles, interpreted as being controlled by an increasing contribution of sulfur-rich, oxidizing magmatic fluids into the hydrothermal system [Chiodini et al., 2002]. The relevant role of the hydrothermal system during the unrest period is underlined by LP seismic events in summer 2001, depicting a source beneath the caldera center at depths of 1–2 km. Geochemical variations after 1997–1999 suggest that flow in the hydrothermal system is unsteady, possibly related to input of hot fluids from the deeper magmatic system [Caliro et al., 2004; Gottsmann et al., 2007].

3.1.2.6 Changbaishan

Changbaishan Tianchi (known as Baitoushan in Korea) stratovolcano lies at the border between China and North Korea, about 1200 km west of the west Pacific subduction zone. The west Pacific plate subducts at ~20° underneath the Eurasian plate along the Japan arc, reaching the mantle transition zone at a depth of ~600 km, where the slab becomes stagnant. Deep dehydration of the slab and convective circulation in the mantle wedge cause upwelling of high-temperature asthenospheric materials, explaining the otherwise unexpected intraplate location of Changbaishan [Zhao et al., 2009; Wei, 2010; Kuritani et al., 2011; Xu et al., 2012; Tang et al., 2014]. A 5 km wide summit caldera (Tianchi) is present on the 2744 m high Changbaishan edifice; the caldera formed during the “Millennium” eruption, which spewed ~100 km³ of peralkaline rhyolite around 939 Common Era [Yin et al., 2012; Yang et al., 2014; Sigl et al., 2015]. Several historical eruptions have been recorded at Changbaishan since the fifteenth century (1413, 1597, 1668, and 1702), although there has been no eruption since 1903 (Figure 11) [Ji et al., 2010].

Prior to 1999, Changbaishan was irregularly and inadequately monitored. Limited data then available hinted at the occurrence of short-lived increased seismicity during 1985–1987, followed by relative quiescence [Wei et al., 2013]. Continuous monitoring of Changbaishan since 1999 has produced a 12 year long time series of activity. During the “inactive period” from 1999 to 2002, there were approximately seven seismic events per month, considered as the background for the region, without significant deformation and with moderate variations the CO₂ content from fumaroles [Xu et al., 2012]. In the 2002–2006 period, the number of earthquakes increased to 72 per month, accompanied by earthquake swarms and harmonic events. The seismicity peaked in November 2003, with 243 seismic events clustering beneath the NE and SW parts of the caldera, at a depth of ~5 km [Wu et al., 2005, 2007; Xu et al., 2012]. A maximum uplift of 68 mm between 2002 and 2005 was detected on the northern caldera rim, related to the addition of magma in a reservoir at 2 to 6 km of depth below the summit, in agreement with the location of earthquake swarms [Xu et al., 2012; Wei et al., 2013; GVP report, 08/2010 (BGVN 35:08)]. The contents of the major gases, as CO₂, H₂, He, and N₂/O₂ components and isotopic ³He/⁴He, were significantly higher during 2002–2006. After 2006, the unrest reversed with a total subsidence of ~12 mm [GVP report, 08/2011 (BGVN 36:08)] but, between 2010 and 2011, an increase in the water temperature of the Julong hot spring was measured (northern portion of the volcano) [Xu et al., 2012].

In synthesis, even if poorly monitored before 1999, the activity of Changbaishan suggests a single, major episode of unrest, with the coupling in the rate of all the monitoring indicators.

3.2 Unrest With Eruptions

3.2.1 Mafic Calderas: Representative Types

3.2.1.1 Fernandina, Cerro Azul, and Sierra Negra

The Galapagos Archipelago lies in the eastern Pacific, 200–300 km south of the Galapagos spreading ridge, separating the Nazca (to the south) and Cocos plates. The volcanic islands of the archipelago are the most recent subaerial expression of the Galapagos hot spot, constructed mainly upon a broad volcanic platform overlying young (<10 Ma) oceanic lithosphere. Fernandina and Isabela, the western Galapagos Islands, are characterized by the construction of seven simultaneously active volcanoes spaced a few tens of kilometers above the inferred plume head (Figure 12) [Feighner and Richards, 1994; Sinton et al., 1996; Werner et al., 2003].

Fernandina, ~30 km wide and 1476 m high, is the westernmost volcano with a well-developed circumferential and radial fissure system. The elliptical summit caldera, NW-SE elongated, is ~900 m deep at its maximum, with walls sloping inward at 30–50°. The caldera results from the coalescence of at least three NW-SE aligned collapses, and at times each was completely filled. The present caldera is affected by the last important collapse of the SE portion, reaching 350 m, in June 1968. A flank eruption on 21 May 1968, minor earthquakes (M_b 3.9–4.6) through the first week of June and a large hydromagmatic explosion from the caldera wall occurred prior to this collapse, which started on 12 June. The combined volume of the two eruptions is <0.2 km³, while the collapse volume is 1–2 km³. No evidence of submarine eruptions during this event has been found, and magma was probably withdrawn from the reservoir through lateral intrusions that did not reach the surface [Simkin and Howard, 1970]. Fernandina is among the most active volcanoes in the world: the post-1968 collapse activity includes eruptions in 1982, 1984, 1988, 1991, 1995, 2005, and 2009. The last eruptions, from fissures that are oriented both radially and circumferentially, have initiated with the intrusion of subhorizontal sills from a shallow magma reservoir [Chadwick et al., 2010; Bagnardi and Amelung, 2012; Bagnardi et al., 2013].

During 14–16 September 1988, a large intracaldera avalanche and an eruption of tephra and lava produced the most profound changes within the caldera since its collapse in 1968. On 14 September 1988, a M_b 4.6 earthquake preceded a 0.9 km³ debris avalanche generated by the failure of a fault-bounded segment of the east caldera wall, ~2 km long and 300 m wide. An eruption (VEI 2) began in the caldera within about 1–2 h of the earthquake, producing a plume and lava flows, ending on 16 September. Most of the eruptive activity was from vents on the caldera floor near the base of the new avalanche scar. The most likely sequence of events seems to be that a dike intruded upward into the east caldera wall dislocated the unstable wall block and triggered the avalanche; the avalanche exposed the newly emplaced dike and initiated the eruption. The cause of the earthquakes is unknown. Between February 2003 and September 2004, LOS (satellite line of sight) displacements (<5 cm) mostly occurred within the caldera rim (Figure 13). Starting from the end of September 2004 and during the beginning of 2005, the deformation propagated in an area covering the entire volcano summit [Bagnardi and Amelung, 2012]. On the morning of 13 May 2005, a circumferential eruptive fissure, formed by a set of five right-stepping en echelon segments, opened on the southern summit plateau of Fernandina. The 2005 circumferential fissure eruption (VEI 2) was fed by a concave shell that is steeply dipping (~45–60°) toward the caldera at the surface and gently dipping (~12–14°) at depth, where it connects to the horizontal subcaldera sill. A flat-topped magma reservoir is postulated at 1.1 km below sea level (bsl), and an oblate-spheroid cavity at ~4.9 km bsl [Chadwick et al., 1991, 2010; Bagnardi and Amelung, 2012]. After the 2005 eruption, the general trend of reinflation observed from the end of 2005 throughout 2006 was interrupted during a seismic crisis offshore east of Fernandina, at a distance of 35 to 60 km, culminating in late August 2007 with M = 3.8 to 5.4 earthquakes: in this period, interferograms showed a rapid negative LOS displacement (movement away from the satellite) of the caldera and of the summit area. After this minor episode of deflation, the caldera uplifted progressively from 2008 to April 2009. During the night between 10 and 11 April 2009, a new eruption (VEI 2) started from radial fissures on the SW flank of Fernandina [Bagnardi and Amelung, 2012]. As during the previous 2005 eruption, the caldera and a large portion of the island show coeruptive negative LOS displacement. The subsiding area outside the caldera is similar to the one displaced during previous episodes as for shape and gradient variations of the interferometric fringes. The early post-2009 eruption interval is characterized by rapid positive LOS displacement only within the caldera. After 5 months from the end of the eruption, beginning in October 2009, the InSAR data record edifice-wide positive LOS displacement. This pattern of deformation persists for more than 7 months until May 2010 [Bagnardi and Amelung, 2012].

In synthesis, the recent eruptive history of Fernandina is characterized by multiple eruptions, with consistent and linear preeruptive behaviors. Deformation increased gradually and was associated, immediately before the eruptions, with short-term seismicity, followed by coeruptive deflation.

Similar preeruptive and eruptive behavior, characterized by long-term inflation, short-term seismicity, and eruption, is shown by the nearby Cerro Azul. Cerro Azul is an active basaltic shield on the SW end of Isabela. The steep upper flanks are surmounted by a wide flat summit rim and a 450 m deep caldera. The nesting indicates several stages of caldera collapse, with different centers of collapse, followed by partial refilling. Historical activity has occurred within the caldera, on the summit rim, and on the flanks, with an average of one eruption every 6.6 years. Ten eruptions have been witnessed between 1932 and 1998. The largest recorded eruption (VEI 3) was from an unknown vent on the summit in 1943. The most recent activity was characterized by two eruptions in 1998 and 2008 (both VEI 1), lasting 51 and 20 days, respectively [Naumann and Geist, 2000, and references therein]. The 3 months immediately after the 1998 eruption showed a rapid inflation (5–10 cm) at the volcano summit, consistent with a source at a depth of ~4 km, and overall lack of seismicity [GVP report, 09/1998 (BGVN 23:09)]. Between 2000 and 2008 inflation continued near the summit, reaching about 30 cm, but with the onset of the 2008 eruption, deflation at the summit was recorded [Baker and Amelung, 2009]. The 2008 eruption from the SE flank lasted from 29 May to 11 June. The first phase involved a rapid emission of lava along tangential or radial eruptive fissures, accompanied by the largest and shallow earthquake, M ~ 3.7, in the northern caldera. The second phase emitted lavas from a separate radial fissure at a lower elevation [Baker and Amelung, 2009; Baker et al., 2011; GVP report, 05/2008 (BGVN 33:05)].

Sierra Negra, on Isabela, is the most voluminous volcano in the western Galapagos. The emerged portion is 60 × 40 km wide and 1140 m high. The summit caldera is larger (7 × 10 km) but shallower (110 m deep) than the other calderas in the western Galapagos, because of resurgent uplift of several tens of meters in the southern part. More than 10 historical eruptions have occurred on Sierra Negra since 1813. Most of the recent eruptions have produced lavas on the north flank, while intracaldera lava flows are older, estimated at 1000–3000 years.

InSAR results from three different intervals between 1992 and 1999 showed that the caldera floor inflated by 2.7 m (Figure 14). From 1992 to 1997, the inflation pattern was nearly axisymmetric with regard to the caldera center. This was modeled as due to intrusion of magma into a sill beneath the caldera, at 2 km of depth. Between 1997 and 1998 the maximum uplift, ~1.2 m of slip along a steeply south dipping normal fault, was centered on the southern limb of a preexisting intracaldera fault system. The focus of inflation at Sierra Negra shifted back to the center of the caldera between 1998 and 1999, again interpreted as magma filling a subcaldera sill [Amelung et al., 2000; Jonsson et al., 2005]. GPS data during 2000–2003 showed a deceleration in the uplift rate, followed by a subsidence of 9 cm/yr [Geist et al., 2006]. The deflationary source was modeled as a contracting sill at a depth of 2.1 km, similar to the inflationary source of the 1990s. In April 2003, deformation of the caldera floor changed from deflation to inflation. The rate of inflation gradually increased throughout 2004 and into 2005, for a total uplift of 89 cm, accompanied by 69 cm of horizontal extension across the caldera. Elastic inflation was interrupted by an episode of inelastic trapdoor faulting marked by a M_b 4.6 earthquake on 16 April 2005; this was responsible for 84 cm of uplift, on the southern part of the caldera, and contraction of the cross-caldera line by 26 cm. Modeling results for this trapdoor faulting suggest slip on a high-angle inward dipping reverse fault, extending from the surface down to the sill at 2.2 km depth. The rate of inflation was not affected by this faulting event, approaching 1 cm/d. Parts of the caldera uplifted 1.22 m between the faulting event and the start of the eruption on 22 October 2005. The (VEI 3) eruption, responsible for a steam and ash plume 13 km high, was characterized by the opening of a fissure inside the north rim of the caldera, on the opposite side of the activation of the M = 4.6 event, with a total eruptive volume of ~150 × 10⁶ m³. No transient deformation occurred in the hours or days before the eruption. During the 8 days of the eruption, geodetic data show that the caldera floor deflated ~5 m, and the volcano contracted horizontally ~6 m [Geist et al., 2008]. Overall, from 1 April 2003 to 22 October 2005, the preeruptive maximum uplift reached ~2.2 m. Horizontal extension across the caldera amounted to 97 cm between 16 April 2005 and the eruption, and a total of 1.4 m since 1 April 2003. This brings the total amount of uplift since 1992 to nearly 5 m, the largest precursory inflation ever recorded at a basaltic caldera. This extraordinary uplift was accommodated in part by repeated trapdoor faulting, with feedbacks between inflation, faulting, and eruption, demonstrating that faulting above an intruding magma body can relieve accumulated strain and effectively postpone eruption. The trapdoor faulting relieved the pressure within the sill but caused compression to the south of the reverse fault. This suggests that while the faulting provides a mechanism for the sill to thicken and postpone eruptions, it also prevents the sill from growing to the south [Chadwick et al., 2006; Jonsson, 2009].

Despite the similar setting and conditions to the other Galapagos calderas (Fernandina and Cerro Azul), the activity of Sierra Negra in the last decades has been affected by a more complex behavior, characterized by a single eruption with a longer and composite unrest and by larger displacements and erupted volumes; in addition, the most uplifted area during the unrest did not coincide with the erupting area. We anticipate that the complexity of Sierra Negra makes it an end-member type among the unrest types of mafic calderas.

3.2.1.2 Kilauea

Kilauea is the youngest and currently most active of five subaerial volcanoes on the island of Hawaii, above a hot spot alive for >70 Ma. The basaltic shield of Kilauea is 1100 m high, and on the summit it hosts a 3.5 × 5 km wide caldera, with a topographic expression of nearly a hundred meters. The most active part of the caldera is the subcircular Halemaumau pit crater, ~1 km wide, with widespread hydrothermal activity. The shallow magmatic system responsible for the activity of Kilauea consists of two long-term reservoirs, at ~3 km beneath the south caldera and at 1–2 km beneath the caldera center (Figure 15) [Poland et al., 2013]. Two volcanic rift zones depart from the caldera: the SW and the more active East Rift, with a kinked geometry. These rift zones feed dike-induced eruptive fissures 5–10 km long and with opening of a few meters per day, as, for example, in 1983. In the last century, Kilauea has erupted a few dozen times, mostly since 1952. Accurate leveling on the summit was begun in 1919–1920, since when almost 5 m of subsidence was recorded, mostly caused by a phreatic eruption in 1924 and the ongoing eruption, which began in 1983. The largest summit subsidence occurred from April to June 1924, when magma moved from a shallow summit reservoir to the East Rift. During the subsidence, steam explosions occurred from Halemaumau, causing its widening. After 1924, summit tilt changes were slow and unrelated to a few small summit eruptions or to a shallow earthquake swarm in May 1938. In the spring of 1950, an increase in the rate of shallow earthquakes was concurrent with uplift of the summit region.

Since 1950 a now familiar displacement pattern has recurred in the summit area. During a period between successive rift eruptions, which may last from a few weeks to a few years, the caldera floor inflates of several millimeters per day. Eventually, the uplift is interrupted by a sudden subsidence, which may coincide with a rift zone intrusion and, possibly, eruption. Within a few hours the caldera floor may drop several tens to hundreds of millimeters. This pattern of vertical and horizontal displacement was explained by pressurization of a vertically elongated ellipsoid [Dvorak and Dzurisin, 1997]. Between November 1975 and January 2008, a significant mass increase occurred beneath Halemaumau Crater. Surprisingly, there was no sustained uplift accompanying the mass accumulation, probably because magma accumulated within a network of interconnected cracks created in response to the 1975 M 7.2 south flank earthquake. Such a refilling represents a gradual recovery from that earthquake. Kilauea has been erupting more or less continuously since 1983 from the Pu‘u‘Ō‘ō and Kūpaianaha vents in the East Rift Zone (ERZ) (mainly VEI 1). With the exception of three brief inflationary periods associated with changes in vent geometry at Pu‘u‘Ō‘ō, the volcano summit subsided continuously between 1983 and 2003 and then it began uplifting [Montgomery-Brown et al., 2010, and references therein]. A new eruptive vent opened at the summit in 2008, probably tapping the reservoir that had been accumulating magma since the 1975 earthquake [Johnson et al., 2010a]. On a longer term, between 1952 and 1991, a relatively constant magma supply rate of 0.11–0.18 km³/yr has been calculated for the shallow magmatic system of Kilauea. However, CO₂ and SO₂ emissions, as well as seismicity and geodetic data, show that during 2003–2007 the supply of magma to Kilauea at least doubled. By the end of 2008, geological, geophysical, and geochemical data indicated a return to pre-2003 rates of magma supply. The transient increase in magma supply to Kilauea was likely caused by a pulse of magma from the mantle that resulted in changes in volcanic activity [Poland et al., 2012]. Geodetic relations between caldera subsidence and magma intrusion along the East Rift of Kilauea or slip of the southern flank have been recently highlighted [Owen et al., 2000; Montgomery-Brown et al., 2010].

Over the last century, Kilauea has erupted a few dozen times, mostly since 1952. Several cycles of deformation, seismic, and gas activities throughout this period occurred, always showing an overall correlation in the rates of the indicators.

3.2.1.3 Okmok

Okmok is a basaltic shield volcano in the central Aleutian Volcanic Arc (Alaska). The arc consists of >40 active volcanoes representing the surface magmatic expression of the subduction of the Pacific plate beneath the North American plate, at a rate of 7–8 cm/yr [Lu and Dzurisin, 2014, and references therein]. Major explosive eruptions of Okmok at ~12 and ~2.05 ka formed two overlapping summit calderas. The younger caldera rim is well preserved and ~10 km wide; the only topographic remnants of its predecessor are two arcuate ridges about 1.5 km north and east of the younger rim. Subsequent eruptions produced small cones and lava flows, including several historically active vents within the younger caldera (Figure 16) [Lu et al., 2005; Lu and Dzurisin, 2014, and references therein]. The Cone D on the east central caldera floor is ~1000 years old, and its deposits suggest highly explosive interactions between magma, groundwater, and surface water. The Cone A on the SW edge of the caldera floor is the youngest, formed almost entirely during the twentieth century [Lu et al., 2010, and references therein]. Okmok has been one of the most active calderas in North America, with eruptions every 10–30 years [Johnson et al., 2010b]. From 1943 to 1997, all of Okmok's eruptions originated from Cone A.

The two most recent eruptions of Okmok, in 1997 and 2008 (VEI 3 and VEI 4, respectively), are much better documented [Larsen et al., 2015]. The 1997 eruption produced abundant ash emissions and mafic lava flows from Cone A; the 2008 eruption formed new vents near Cone D [Lu et al., 2000; Biggs et al., 2010, and references therein]. The inflation/deflation rate at Okmok was not steady during 1992–2008, which might indicate that the magma supply from a deep production zone varied with time. The average surface uplift rate decreased from about 10 cm/yr during 1992–1993 to 2–3 cm/yr during 1993–1995, and the surface subsided at an average rate of 1–2 cm/yr during 1995–1996. After such a stasis of deformation, the 1997 eruption began, lasting for ~2 months and producing a lava flow of 1.5 × 10⁸ m³ [Mann et al., 2002; Lu et al., 2003; Biggs et al., 2010; Lu and Dzurisin, 2014] (Alaska Volcano Observatory (AVO) report 2001, http://www.avo.alaska.edu). After the eruption, from September 1997 to September 2000, 20 cm of uplift were observed within the caldera [Mann et al., 2002]. During May 2001 an earthquake swarm with M_L 2.0–3.6 occurred without any volcanic activity or thermal anomalies (AVO report 2001, http://www.avo.alaska.edu). The inflation lasted until the spring-summer of 2004 [Fournier et al., 2009; Biggs et al., 2010]. Subsequently, from 2005 to early 2008 periods of stasis in the deformation and slight deflation alternated. The second eruption started on 12 July 2008; the network recorded no significant precursory seismicity until 5 h prior to the eruption, when small detectable earthquakes occurred at a rate of 5–15 events per hour [Pesicek et al., 2012; Johnson et al., 2010b].

Summarizing, in the last decades Okmok erupted in 1997 and 2008. In both cases the preeruptive deformation behavior was similar: an increase in the deformation rate followed by a period of stasis, with a final eruption. The absence of significant inflation immediately before the eruptions, during 1995–1997 and 2004–2007, suggests a critical state of pressurization for the reservoir, in which the surrounding country rock is strong enough to temporarily slow or halt the magma supply from depth. Exsolution of magmatic volatiles during these periods and resulting pressurization in the upper part of the reservoir might have triggered the 1997 and 2008 eruptions [Lu and Dzurisin, 2014].

3.2.1.4 Recent Collapsing Calderas: Miyakejima and Dolomieu (Piton de la Fournaise)

Miyakejima is a basaltic-andesitic stratovolcano with a summit caldera lying in the northern Izu-Bonin Arc. It is associated with the subduction of the Pacific plate beneath the Philippine Sea plate, on which the island of Miyakejima is located. In turn, the Philippine Sea plate moves northward and subducts beneath the Eurasian plate along the Sagami-Nankai trough, about 100 km NW of Miyakejima. The subaerial portion of Miyakejima is ~8 km wide and 0.7 km high, with at least three nested summit calderas. The recent eruptive activity includes >10 radial fissure eruptions from the sixteenth century to 1983 [Tsukui et al., 2005]; the last occurred on the SW flank on 4 October 1983 and formed a 4.5 km long eruptive fissure, from which 7 × 10⁷ m³ Dense Rock Equivalent (DRE) of scoria and lavas were emitted (Figure 17) [Tsukui and Suzuki, 1998].

On 26 June 2000, earthquakes occurred at depth <5 km bsl, initially beneath the SW flank near the summit and gradually migrated west of the island, where a submarine eruption occurred the next morning [Uhira et al., 2005]. Almost simultaneous with the onset of the earthquake swarms, ground deformation toward east and upward was observed in the SE part of the volcano at 18:00 on 26 June 2000. After 21:30 on 26 June 2000, the displacements turned from eastward to westward. Three hours later, the displacement rates increased gradually in the western part of Miyakejima, as the seismicity migrated and approached the coast and reached a climax with the submarine eruption [Meilano et al., 2003]. From 27 June, earthquakes migrated farther to the NW, clustering along a NW-SE direction and ascending slowly with time. The seismic swarm continued to late August, consisting of 7000 shocks with magnitude M ≥ 3, and five M ≥ 6 shocks [Toda et al., 2002; Uhira et al., 2005]. During this seismic crisis, on 8 July a caldera, 1.6 km wide and 450 m deep, developed at the summit of Miyakejima with a VEI 3 phreatic eruption. Growth of the caldera took from 8 July to the middle of August, with a mean rate of 1.4 × 10⁷ m³/d and a final collapsed volume of 6 × 10⁸ m³. The total subsidence of the caldera floor was 1.6–2.1 km [Kumagai et al., 2001; Geshi et al., 2002]. This sequence of events is compatible with a laterally propagating dyke, which partially emptied the magma chamber and formed the caldera [Toda et al., 2002]. In particular, at 18:30 LT on 26 June, the magma began to ascend from a dyke-shaped magma chamber and intrude below the SW flank. At ~21:00 LT, a relatively large volume of magma began to intrude laterally toward the NW as a dyke beneath the west coast. This intrusion caused a discharge of magma from the chamber. The NW propagation of the dyke and the contraction of the chamber continued thereafter. The discharge of magma from the chamber probably starved the first intrusion that had been ascending toward the SW flank of the island, resulting in the collapse-caldera formation after the initial stage. The 2000 eruptive activity differed from other recent activity in the huge subsurface discharge of magma toward NW [Ueda et al., 2005]. The SO₂ emission reached 10⁵ t/d in October–November 2000. Degassing from the magma reservoir deflated the edifice until 2004, when the deformation turned to slow inflation, possibly because of the reinjection of magma into the reservoir [Kazahaya et al., 2004].

A very similar sequence of events was observed at Piton de la Fournaise, Réunion, in 2007. The island of La Réunion (SW Indian Ocean) consists of a hot spot-induced flattened volcanic edifice, 230 km wide at the level of the surrounding seafloor and 7 to 8 km high. It consists of two subaerial volcanic massifs: Piton des Neiges and Piton de La Fournaise, developed in last 2.1 Ma [Charvis et al., 1999; Gallart et al., 1999]; the latter, on the SE part of the island, is a highly active basaltic shield with four nested U-shaped calderas that open seaward, developed in the last ~150 ka [Staudacher and Allègre, 1993; Gillot et al., 1994; Merle et al., 2010]. The recent activity of Piton de La Fournaise focused on the central cone and along its NE and SE rifts. The cone summit shows two collapsed structures, the currently inactive Bory crater and the Dolomieu crater, which is the locus of numerous summit eruptions. Dolomieu developed by sudden collapses above a 1.5 km wide high-velocity plug located above sea level and corresponding to an intrusive, solidified dyke-and-sill complex. Below, two low-velocity anomalies, from 1 to 0 km asl and below 1 km bsl, may correspond to magma storage systems [Prôno et al., 2009].

In the last decades, two periods of frequent eruptions separated by 6 years of inactivity occurred at Piton de la Fournaise: 1972–1992 and 1998–2007. During 1972–1992 and 1998–2000, shorter-term precursors anticipated eruptions fed by the progressive drainage of a shallow magma reservoir, recharged only in 1986 and 1998. Conversely, from 2000 to 2007, longer-term precursors preceded cycles of successive eruptions. Each cycle showed summit and near-summit eruptions, ending with a distal eruption on the eastern flank, with the refill of the reservoir by deeper magmas. An example is the December 2002 to September 2003 eruptive cycle, first with vertical dikes feeding summit eruptions and then by dikes along the same path, but subsequently propagating laterally, intruding the rift zones and displacing seaward the eastern flank. The preferential motion of the eastern flank caused by continuous recharging of the shallow reservoir since 2000 favors the distal eruptions toward this flank at the end of a cycle [Peltier et al., 2007, 2009].

The April 2007 caldera formation at Dolomieu was anticipated by continuous inflation from January 2007, followed by sudden, large-scale deflation on 30 March; in this period, two minor eruptions occurred, from Dolomieu (18 February) and on the SE flank, at 1800 m asl (30 March) [Staudacher et al., 2009]. Deflation increased on 1 April, while no magma was erupted, followed on 2 April by a distal eruption (VEI 1), 7 km away on the lower eastern flank of the volcano, at 600 m asl. The summit deflation was followed by an increase of the tremor at the eruption site, suggesting a direct effect of the subsidence on the magma chamber. On 5 April, the summit caldera collapsed, displacing the 1 km wide floor of Dolomieu; the collapse was incremental, with cyclic seismic signals, totaling a subsidence of ~330 m. While the eruptive succession was triggered by a deep magma input, the caldera collapse itself occurred during rapid emptying of the shallow plumbing system feeding the effusive eruption at 600 m asl due to a laterally propagating dike [Michon et al., 2007, 2009].

In synthesis, both Miyakejima and Piton de la Fournaise underwent the incremental development of a ~1 km wide summit caldera following lateral drainage of magma from a shallow reservoir. The incremental behavior observed during collapse at Miyakejima, Dolomieu, and at Fernandina in 1968 results from the friction developed along the preexisting caldera ring faults, as also experimentally confirmed [Michon et al., 2011; Ruch et al., 2012]. Both Miyakejima and Dolomieu inflated after the last eruption, with an increase of deformation and intensification of seismicity only shortly before caldera collapse. The short time span for the precollapse seismicity may be explained by the eruptive frequency of the volcanoes, which, especially at Piton de la Fournaise, maintains an open conduit for the rise of magma until very few kilometers to the summit.

3.2.1.5 Other Mafic Calderas

The eruptive unrest episodes at the above mentioned mafic calderas show an overall recurrent behavior, characterized by the progressive increase in the rates of the indicators, usually coupled deformation and seismicity, before the eruption. Where magma is already shallow, the seismicity may be only a very short term indicator and much shorter than observed inflation. This general behavior of mafic calderas may have minor complications, as at Sierra Negra, or even develop caldera collapse, as at Miyakejima and Dolomieu. However, despite these deviations, the basic regular behavior, testified by the progressive increase in the rate of the indicators prior to the eruption, is usually maintained by most mafic calderas undergoing eruptive unrest. This behavior is common also to other basaltic calderas, briefly summarized below.

Izu-Oshima (Japan) erupted once in 1990, with VEI 2. A M = 6.5 earthquake on 20 February 1990 occurred offshore ~10 km west of the island, with N-S trending aftershocks. On 1 March, a swarm of small earthquakes was recorded near the summit, followed by 10 weak tremor episodes in April. A VEI 2 eruption at Mihara-yama, a scoria cone formed in 1777, occurred on 4 October 1990, ending on the same day. Seismicity started shortly before and continued during the eruption, with 633 events in October, 160 in November and 57 in December [GVP report, 12/1990 (BGVN 15:12)]. Successive, noneruptive unrest episodes occurred in 1993–1999, 2004, 2007, and 2010. On 30 May 1993, continuous volcanic tremor at the active Mihara-yama cone was coeval to an increase in shallow seismic activity below the summit; both the tremor and earthquake swarm stopped by 5 June [GVP report, 05/1993 (BGVN 18:05)]. During 1993–1999, a gradual subsidence of 3.0 ± 0.7 cm/yr within the caldera was associated with weak seismic activity [Furuya, 2005]. In 2004 and 2007 high seismicity was synchronous with inflation of the edifice. In May 2010 inflation preceded an increase in seismicity in July [Furuya, 2005; GVP report, 01/2013 (BGVN 38:01)]: both events probably resulted from magma intrusion. In September, the inflation began declining. However, seismicity continued at a low level, while in the subsequent months it gradually increased to the west offshore of the island. On 9 February 2011, earthquakes again increased in frequency. GPS and strainmeter measurements indicated contraction since January, but the trend reversed to show inflation in October 2011 and seismicity remained at a low level [GVP report, 01/2013 (BGVN 38:01)]. No remarkable activity has been noted since October 2011 (Japan Meteorological Agency (JMA) report OCT20122, http://www.jma.go.jp/jma/indexe.html).

Aoba (Vanuatu) erupted in 1995 and 2005. Both eruptions, with VEI 2, had a similar behavior, with increasing seismicity and gas emissions. Seismic tremor occurred months before the 1995 eruption, while in the 2005 eruption the seismicity was a shorter-term indicator [Rouland et al., 2001; GVP report, 12/2005 (BGVN 30:12)].

Karthala (Comores Archipelago, Indian Ocean) erupted in 1991 (VEI 2), April 2005 (VEI 2), November 2005 (VEI 3), May–June 2006 (VEI 0), and January 2007 (VEI 2). Conversely to the 1991 phreatic eruption, during which ash fallouts were confined to the summit, the highly dynamic eruption of November 2005 caused ash fallouts over the island. However, a similar preeruptive behavior was given by rapid seismicity increasing days/months before the onset of the eruption, without significant ground deformation [Savin et al., 2005; Morin et al., 2009, and references therein].

Katla (Iceland) likely experienced a minor subglacial eruption in 1999. This was preceded by minor unrest with seismicity and inflation, consistent with a point source at 3.5 km depth on the volcano flank. A burst of seismic tremor preceded a jökulhlaup from the glacier covering the volcano on 17–18 July 1999; the subglacial eruption has been inferred to be the cause for the jökulhlaup [Sturkell et al., 2003; GVP report, 09/1999 (BGVN 24:09)]. From 1999 to 2004, GPS measurements along the caldera edge revealed steady inflation of the volcano at a rate of up to 2 cm/yr, together with outward horizontal displacement of far-field stations (>11 km) at ~0.5 cm/yr [GVP report, 11/2011 (BGVN 36:11)]. There was concern after the 2010 eruption of Eyjafjallajokull that an eruption from the Katla caldera might follow, but up to this time of writing, it has not.

Masaya (Nicaragua) has shown repeated eruptions (all VEI 1) in 1993, 1996, 1997, 1998, 1999, 2001, 2003, 2004(?), 2005, 2006, and 2008. This almost constant eruptive activity has been commonly preceded by higher SO₂ flux values and tremor since approximately a month before; there is no report on any preeruptive deformation [Métaxian et al., 1997; Rymer et al., 1998a; GVP report, 11/2011 (BGVN 36:11)].

Poas (Costa Rica) underwent seven recent eruptions in 1991, 1992–1993, 1994, 1996, 2006, 2008, and 2009–2014. These usually were low magnitude (VEI 1), characterized by phreatic explosions, or geyser-like ejections of crater lake water, with the exception of the VEI 2 1994 eruption (a block-and-ash eruption with phreatic and fumarolic activity); the preceding unrest episodes were always characterized by minor uplift (<0.1 cm/yr), usually minor but increasing seismicity (relevant in only two cases, 1994 and 2009–2014) and degassing manifested on average 3 months before the eruption [GVP report, 09/2013 (BGVN 38:09)].

Pacaya (Guatemala), on the edge of the older Amatitlan caldera, has been erupting almost continuously in the last decades. However, in three eruptive periods, between 1990 and 2002 (VEI 3), 2004–2010 (VEI 3), and 2013–2014 (VEI 2), the eruptions were often preceded by increases in seismicity. There is no available information on ground deformation [GVP report, 05/2014 (BGVN 39:05)].

Planchón-Peteroa (Argentina-Chile) erupted in 1991 (VEI 2), 1998 (VEI 1), 2010 (VEI 2), and 2011 (VEI 3). Preeruptive (occurring at least 15 days before the onset of eruptions) seismicity was poor in 1991 and almost absent in 1998, while in 2010 and 2011 it was intense and with continuous tremor, in general, consistently with the size of the impending eruption. There is no available information on ground deformation [GVP report, 11/2013 (BGVN 38:11)].

3.2.2 Felsic Calderas: Representative Types

3.2.2.1 Rabaul

Rabaul caldera, Papua New Guinea, is associated with the subduction of the Solomon plate beneath the Bismarck plate. Rabaul consists of a pair of nested collapse structures. While the larger appears inactive, the inner collapse is easily identified by the focused seismicity; this denotes a strongly elongated caldera formed by high-angle outward dipping faults above a magma chamber at 3–6 km of depth [Mori and McKee, 1987; Finlayson et al., 2003; Bai and Greenhalgh, 2005]. Rabaul has experienced several caldera-forming eruptions. After the last caldera-forming eruption, between CE 667 and 699 [McKee et al., 2015], more than eight intracaldera eruptions have occurred, the last three in 1878, 1937–1941, and between 1994 and the present. These last three eruptions have been characterized by the simultaneous activity of the Tavurvur and Vulcan basalt-dacite eruptive centers, which lie along the inner, active caldera, but on opposite sides. Their activity seems to be related to the injection of hotter mafic magma within silicic melts (Figure 18) [Johnson et al., 1995; Nairn et al., 1995; Roggensack et al., 1996; Bouvet de Maisonneuve et al., 2014a, 2014b].

The recent history of Rabaul has been marked by restless behavior, with several unrest episodes, culminating in an eruption sequence starting in 1994. Following two M 8 tectonic earthquakes in the Solomon Sea in 1971, Rabaul underwent progressive uplift and tilting, concentrated in its central part; the uplift was accompanied by shallow seismicity (<3 km) along the inner ring structure, significantly greater than before 1971. Gravity changes also reflected the uplift of the caldera. Between 1971 and 1983, several thousand earthquakes were accompanied by a cumulative uplift of 1 m near the caldera center. These earthquakes along the inner ring fault have been interpreted as evidence that magmatic pressure was beginning to build up and that small, mafic magma injections were taking place. In 1983, both the seismicity and uplift rate dramatically increased. This increase may have been related to a M 7.6 regional earthquake occurring 200 km east of Rabaul in March 1983. The acceleration of the unrest lasted until July 1985 and was characterized by several tens of thousands of earthquakes and an uplift of nearly 1 m. The 1983–1985 unrest resulted from a major injection of mafic magma; considerable magmatic pressure was built up as a result of magma mixing, but this was still insufficient to cause an eruption. From August 1985 to April 1992 the seismicity fluctuated at levels comparable to those of the 1971–1983 period, even though the uplift rate clearly decreased, with ~35 cm in nearly 7 years. From May 1992 to September 1994 the uplift rate increased again, becoming comparable to that observed between 1971 and 1983; the net uplift in the central part of the caldera was approximately 20 cm. At the same time, seismicity moderately increased relative to the 1985–1992 period; interestingly, part of the observed seismicity is interpreted to result from the injection of mafic magma into a shallow felsic reservoir below the central part of the caldera. A burst of seismicity, probably including also low-frequency earthquakes, occurred in late August 1994. However, no seismicity was recorded from 29 August to 18 September, when a M 5.1 earthquake occurred below the eastern part of the caldera, at a depth of 1.2 km. After only 27 h of sustained seismicity, 12 of which included low-frequency events, the eruption began on 19 September. Tavurvur, on the east rim, erupted first, through one vent; here the uplift immediately preceding the eruption was 1–2 m. Approximately 1–1.5 h later Vulcan, following a localized uplift of 6 m a few hours before the eruption, also started to erupt, from at least four vents. Several tsunamis were generated during the onset of the eruption. Soon after the eruption onset, geodetic data indicated a deflation of the caldera. The initial part of the VEI 4 eruption was the most violent, generating an ash cloud reaching ~20 km of altitude. The Surtseyan-type of eruption from Vulcan was the more powerful and included a brief phase of strong Plinian activity soon after its onset; however, Vulcan's eruption ended on 2 October. The eruption at Tavurvur, after peaking during the first 5 days of activity, lasted several months, even though with progressive decline. At the end of October, the subsidence of the caldera reached at least 1 m in the central part, and 20–30 cm along the edges. Seismic activity along the caldera rim progressively decreased from the beginning of October to the end of November. The Tavurvur eruption ended in mid-April 1995. From November 1995 to July 2010, at least five other explosive eruptive events, with 1 < VEI < 4, occurred [Williams, 1995; Johnson et al., 1995, 2010c].

In synthesis, the most recent unrest episodes at Rabaul have been characterized by an overall coupling between the increase in volcanic seismicity and deformation. However, we note also two remarkable behaviors of this restless erupting caldera. In 1971 and 1983, the unrest episodes, or their intensification, occurred immediately after major regional earthquakes, suggesting interactions between tectonic events and a prepared magmatic system. However, while Rabaul responded to a M = 7 earthquake at a distance of 180 km with a pronounced earthquake swarm, it did not produce any detectable activity in response to a second M = 7 earthquake 2 months later, at a distance of only 60 km [Mori et al., 1989]. The second is that high seismicity and uplift between 1983 and 1985 was not followed immediately by any eruption, while moderate seismicity and uplift occurring between 1992 and 1994 did culminate in an eruption. This behavior may be understood only by considering the second unrest and eventual eruption as a continuation and dependent on the first. The first primed the system, and the second set off the eruption.

3.2.2.2 Pinatubo

A low-end VEI 6 eruption of Pinatubo on 15 June 1991 led within hours to formation of a 2.5 km wide caldera. The 1991 caldera is just the latest of the multiple calderas that have formed and filled during Pinatubo's history. The largest prehistoric caldera, ~5 km in diameter, was probably formed during the largest known eruption, ~81 ka [Newhall et al., 1996; Ku et al., 2008]. Collapse in 1991 was apparently piecemeal, and one block on the NE side of the caldera that did not collapse now exposes old caldera fill (moat) sediments. The rest of the 1991 caldera walls expose remnants of lava domes that at various times filled older calderas and rebuilt the Pinatubo edifice. Abundant basaltic blebs in hybrid andesite of these domes suggested a recurring process—basalt intrusion into dacite magma, magma mixing, and dome extrusion often but perhaps not always followed by large explosive eruptions. Similar domes form after large explosive eruptions.

Unrest at Pinatubo may have begun within weeks after the 16 July 1990 M 7.8 earthquake along the Philippine Fault, about 100 km ENE of Pinatubo (Figure 19). There were reports of landslides and steaming within weeks after the earthquake, and some M 4–5 earthquakes were located in the general vicinity of Pinatubo but with only the low precision of the worldwide seismic network. No local monitoring existed. Pinatubo was then forgotten until mid-March 1991, when new local earthquakes began, and 2 April 1991, when a NE trending fissure opened across the north flank of the volcano, phreatic explosions formed a chain of craters, and strong steaming began from vents at the end of the fissure nearest the volcano. Local seismic monitoring was started within days—first using a tripartite, nontelemetered network [Sabit et al., 1996] and then using seven radio-telemetered stations around the entire volcano [Lockhart et al., 1996; Murray et al., 1996]. Early in April, water in the Maraunot River 1.5 km from the summit was unusually acidic (pH 1.95) [Campita et al., 1991; Sabit et al., 1996], suggesting expulsion of acidic fluid from the preexisting, known geothermal system. Seismicity in April and most of May was strongest in the first weeks and was mostly in a cluster of hypocenters ~5 km NW of the summit, along a preexisting regional fault (the Iba-Maraunot Fault). In retrospect, this was an example of distal volcano-tectonic seismicity triggered by inflation of a magma or hydrothermal source beneath Pinatubo and consequent compression of a confined aquifer ~5 km NW of the summit. For nearly the entire 2 month precursory period, seismicity fluctuated without obvious trend, was dominated by volcano-tectonic (VT) events, and was unremarkable. Nothing in the seismicity of April and most of May, or even until 14 June, two full days into the precursory VEI 3 eruptions, would have been diagnostic of the impending VEI 6 eruption. Successful forecast of the VEI 6 eruption was based wholly on the prevalence of VEI 6 events in the geologic record.

On 26–28 May and again on 3 June, the network recorded 30–35 km deep long-period (DLP) earthquakes that were not recognized as such at the time [White, 1996]. In retrospect, these can be interpreted as evidence of magma resupply at depth. Did magma movement begin at that depth and moved upward? Or did prior, aseismic ascent of magma sometime before these DLPs induce further magma ascent from depth? The latter seems more likely, since (a) static compression of +1 bar had been “squeezing” the crust beneath Pinatubo since the earthquake of 1990 (R. Stein in Bautista et al. [1996]) and (b) there is independent evidence from olivine diffusion rims that fresh basalt mixed with residual dacite at least as early as March 1991 (M. Coombs, personal communication, 2002). Unfortunately, seismic recording before early May was inadequate to resolve any earlier DLP events.

Monitoring of SO₂ emission was finally started on 13 May, with a measurement of 490 t/d. By 28 May the flux was 5020 t/d, after which it dropped to just 260 t/d on 5 June before recovering and going offscale on 10 June. The overall increase was interpreted as magma ascent, but it could also have reflected diminishing scrubbing of the SO₂ by the groundwater system beneath Pinatubo, i.e., a progressive drying of a “chimney” through the groundwater through which SO₂ gas could escape. The sudden but short-lived near shutoff of SO₂ on 5 June was interpreted as quenching and development of a temporary, impermeable carapace over rising magma, but it could have also reflected temporary flooding of the dry chimney as new fracture permeability developed above the vanguard magma. There was, unfortunately, no time or funding to set up much deformation monitoring before the eruption. This was before the era of GPS and InSAR, and the only instrumental deformation measurements obtained were from two tiltmeters installed high on the cone in late May to early June, and an observation of no macroscopic (tape-measured) extension across the fissure from 1 to 28 May [Ewert et al., 1996]. The tiltmeters did show a few tens of microradians of tilt from 4 to 7 June, but nothing more than what might be expected at proximal tilt sites before any eruption.

Starting on or about 3 June, seismicity shifted from the NW cluster to a new cluster beneath the summit and migrated broadly upward over the next week. Energy release picked up, and on 5 June, the alert level was raised to level 3 (“eruption possible within 2 weeks”) [Punongbayan et al., 1996]. By 7 June there were so many small, shallow events that Alert level 4 was raised (“eruption possible within 24 h”). And, indeed, an eruption did begin that same day, but only extrusion of a hybrid andesite dome and not the expected plinian eruption. Seismicity continued to increase, and on 12 June, a series of ~17 VEI 3 subplinian eruptions began. Initially, magma was a hybrid andesite, but by the afternoon of 13 June it was wholly dacite, and eruptions were becoming smaller and more frequent [Hoblitt et al., 1996; Wolfe and Hoblitt, 1996]. Throughout 12–13 June, seismicity was continuing to increase, and on 14 June there was a dramatic increase in the energy release from shallow LP earthquakes [Harlow et al., 1996]. This last increase was the first and only monitored unrest that was truly diagnostic of a huge impending eruption. Apparently, it was not until 14 June that opening up the conduit and tapping of increasingly gas-rich dacite magma became irreversible, runaway processes. The climactic phase of the eruption began at 1342 h local time on 15 June—Black Saturday. A 500 km diameter umbrella cloud darkened almost all of northern Luzon, and ashfall mixed with rain from Typhoon Yunya (passing near Pinatubo at the same time) caused many roofs to collapse. M 4–5 earthquakes late on 15 June and into 16 June were thought to reflect formation of the caldera, but the few focal mechanisms that could be determined suggest strike-slip motion along the faults that pass through Pinatubo's summit, tectonic readjustment to drawdown of the magma reservoir.

In comparison to unrest at larger calderas, that at Pinatubo was more akin to unrest at stratovolcanoes. Intrusion of fresh mafic magma was buffered for a few months by the residual dacite, but not for years as is often the case at larger calderas. A geothermal system hid SO₂ emission, by dissolving the incoming SO₂ until heat from the SO₂ and more abundant H₂O and CO₂ dried out a chimney through that geothermal system. In the end, magma rose through a single conduit near the summit, and the caldera did not form until near the end of the eruption. One similarity, though, is that the final, irreversible acceleration to climactic eruption did not occur until just 1 day before that climactic eruption began—similar to the 1–2 day final accelerations toward smaller eruptions at Campi Flegrei (1538) and Rabaul (1994). Why did Pinatubo produce a VEI 6 eruption while those of Campi Flegrei, Rabaul, and many other calderas in historical time are smaller, far short of the largest prehistoric eruptions of those systems? Part of the answer may lie in plugging of the conduit at Pinatubo and consequent accumulation of gas far in excess of saturation [Gerlach et al., 1996; Wallace, 2001]. This distinction between plugged versus open conduits and the hypothesis that a plugged conduit and excess volatile accumulation are prerequisite for silicic plinian eruptions is further discussed in Whelley et al. [2014], Winson et al. [2014], and Newhall [2015]. Another part of the answer was suggested by Hoblitt et al. [1996] and Scandone et al. [2007] who, from decreasing repose periods and decreasing microlite contents of 12 June to early 15 June, inferred progressive interconnection or integration of fractures from the magma chamber to the surface, and a correspondingly higher rate of magma ascent. The excess volatile phase in the magma chamber, combined with the connection of fractures to the surface, led to runaway conditions and the plinian eruption that lasted for hours, apparently until the gas-rich top of the magma chamber was exhausted.

3.2.2.3 Iwo-Jima

Frequent but apparently episodic unrest at Iwo-Jima Caldera up through the mid-1980s was discussed by Newhall and Dzurisin [1988, and references therein]. Up until then, there had been ~100 m of uplift (active resurgence) of the caldera floor within the past 400 years [Kaizuka et al., 1985; Kaizuka, 1992; Newhall et al., 1998]. Prominent terraces suggested either episodic uplift or periods of exceptional wave erosion (e.g., during typhoons). Phreatic eruptions occurred every few years or decades, and some post World War II runways were cut by fault displacements of up to 6 m. Digging <1 m into many parts of the island would reveal hot soil (tuff). A comprehensive summary of the geology and unrest of Iwo-Jima was published in a special issue of the Journal of Geography (Tokyo), 1985, titled Geosciences on Iwo-Jima (multiple authors) New details of Iwo-Jima's eruptive history are given by Nagai and Kobayashi [2015].

Unrest has continued to the present, and the nature of the episodicity has become clearer with the advent of GPS, InSAR, and other monitoring [Ukawa et al., 2006] (Figure 20). The general pattern appears to be short periods (weeks) of rapid uplift, often in the order of a meter, followed by several years of partial relaxation of that uplift. The center of uplift is usually near the high point of the resurgent dome and center of the caldera (Motoyama), but it has occasionally been in fault-bounded blocks outboard of the center, in the Chidorigahara area. From 1977 to 1995, long-term subsidence of Motoyama reached 0.54 m and uplift of Chidorigahara exceeded 3 m. Relatively long, slow changes were punctuated by broad, episodic uplift of more than 1 m during volcanic unrest both in 1982 and 2001 [Ukawa et al., 2006]. Based on dominance of vertical changes, the source of episodic uplift is thought to be broader and/or deeper than the 0.10–2.4 km deep source of subsidence [Ukawa et al., 2006]. In August 2006, subsidence was again broken by rapid uplift, reaching 1.2 m by September 2007 [Ozawa et al., 2007]. Long-term net uplift continues at ~25 cm/yr, a remarkable rate matched consistently, to our knowledge, only by uplift of the Yenkahe block of the Siwi Caldera of Vanuatu (150 m in 1000 years) [Chen et al., 1995; see also Erre, 2005; Métrich et al., 2011; Peltier et al., 2012] and, since 2007, remarkable uplift at Laguna del Maule, Chile [Feigl et al., 2014; Singer et al., 2014]. The average rate of inflation at Iwo-Jima is approximately 10⁷ m³/ yr, and that rate can also be taken as a minimum rate of magma supply into its plumbing system. Modeling of the uplift suggests that intrusions reach to within (a few) kilometers of the surface, and drilling for geothermal power in (year) reached fresh syenite at a depth of only 600 m. Partial relaxation of uplift may be the result of bleeding off of pressure within the hydrothermal system after each fresh intrusion and/or degassing of vesiculated, inflated magma. Both processes seem likely. Like many geothermal systems around the world, that of Iwo-Jima seems to increase the sensitivity of Iwo-Jima to disturbances from remote tectonic earthquakes, with repeated, brief increases in seismicity following passage of the surface wave of distant earthquakes [Ukawa et al., 2002]. Although small phreatic eruptions continue, they are inconsequential relative to the magnitude of ongoing magma resupply and resurgence.

Uplift at Iwo-Jima gained prominence in the study of coastal processes and terrace formation [Kaizuka et al., 1985; Shigemura, 1986; Kaizuka, 1992] but was brought to the attention of volcanologists by unrest at Long Valley, Campi Flegrei, and Rabaul in the early 1980s. Given the extremely high rates of deformation, an early concern was that a new caldera-forming eruption might be at hand. However, as understanding of plugged and open systems has evolved since then, it now appears that Iwo-Jima is an example of the openly degassing end-member of calderas. The most recent known magmatic eruption occurred just shortly after the top of the resurgent dome rose through and above wave levels, ~400 years ago. The remarkable episodic magma supply and uplift since that time seems not to have triggered any magmatic eruption. The reason may be that nearly all of the incoming gas can escape freely (Iwo-Jima means “sulfur island”), preventing any buildup and magmatic eruption. In this sense, Iwo-Jima is the polar opposite of Pinatubo, at least within historical time.

3.2.2.4 Aso

Aso caldera lies along the western continuation of the transtensive Median Tectonic Line, in central Kyushu Island, SW Japan. The Median Tectonic Line is the surface expression of the oblique convergence of the Philippine Sea plate subducting below the Eurasian plate of west Japan [Kamata and Kodama, 1999; Takayama and Yoshida, 2007]. Aso has a 18 × 25 km N-S elongated caldera, formed after four major explosive eruptions with a total volume >200 km³, from 270 to 90 ka. Postcaldera activity, consisting of >17 basaltic to rhyolite cones and domes, followed in the central part. Of these vents, only Nakadake is active and consists of a composite basaltic to andesitic volcano [Miyabuchi, 2011; Miyoshi et al., 2012].

Leveling surveys from 1937 to 2003 show an overall deflation of ~7 cm within the caldera center, interrupted by an uplift of 4 cm in the 1950s. Between 1992 and 2003, the surface of the caldera contracted, subsiding ~2 cm [Sudo and Kong, 2001]. Nakadake had a few phreatomagmatic eruptions between 1989 and 1991, followed by other phreatomagmatic eruptions in 1992 and 1995 and a phreatic one in 1994 (all with VEI 2). These eruptions were preceded by preeruptive seismicity, but in a context of overall deflation. Since 1995, the crater bottom has been occupied by a lake, whose temperature rose from 28° to 82° until 2008 [Ikebe et al., 2008; Miyabuchi et al., 2008]. Seismicity recorded between January 2000 and April 2003 was generally constant, with continuous volcanic tremor every month, in addition to isolated tremor events. The number of tremor events was higher through October 2000, during April 2002, and from August 2002 through March 2003: in 2002, for the first time since 1995, isolated tremor events occurred at a rate of >300 events per day within Nakadake [GVP report, 10/2003 (BGVN 28:10)]. A phreatic eruption (VEI 1) occurred in July 2003. GPS baselines and 2-D strain revealed extension in the caldera in 2003 and compression before and after 2003, until 2010. The seismicity and deformation data suggest that in 2003 a small volume of magma intruded a sill at ~15 km of depth [Unglert et al., 2011]. Between 2003 and 2005, multiple ash emissions, related to a newly ascending magma, occurred from the crater lake [Miyabuchi et al., 2008; Ikebe et al., 2008]. In 2005, an intense seismic activity was recorded, reaching up to 4000 events in March. A more recent unrest episode occurred from April 2011 to September 2012, with mud eruptions, beginning on 6 May 2011, preceded by intense seismic activity, without ground deformation: these eruptions, largely phreatic, usually produced plumes <2 km high [GVP report, 08/2012 (BGVN 37:08)]. The most recent eruption phase, from Nakadake, anticipated by seismic activity, began in November 2014 and continued at least through March 2015 (report JMA, http://www.jma.go.jp/jma, Monthly Volcanic Activity Report (MVAR) December 2014).

In synthesis, in the last decades, Aso produced several minor largely phreatic eruptions (VEI ≤ 2) in 1989, 1992, 1994, 2003–2005, 2011–2012, and 2014–2015. Seismicity has been a precursor for every eruption, in some cases since months before (as in the 2003 and 2011 eruptions), in others just days before (as in the 2005 eruption). However, all the eruptions, with the possible exception of 2003, occurred in a context of minor deflation. The occurrence of several phreatic explosions, as well as the complex dynamics of the crater lake, suggests that a shallow hydrothermal system plays an important role in the preeruptive dynamics of Aso.

3.2.2.5 Toya (Mount Usu) and Aira (Sakurajima)

Toya and Aira (Japan) are calderas both characterized by the recent activity of a stratovolcano (Mount Usu and Sakurajima, respectively) on their southern rim. Toya Caldera lies in western Hokkaido, NE Japan, at the V-shaped intersection between the Kurile and Japan arcs, overlying the Pacific plate, subducting in a NW direction at ~10 cm/yr. The caldera, formed by collapse at 115–112 ka, is a nearly circular depression 10 km wide filled with Lake Toya. After caldera formation, Nakajima and Usu volcanoes formed at the center and the southern edge of the caldera, respectively. Usu consists of a basaltic edifice with a 1.8 km wide summit caldera and several dacitic lava domes and cryptodomes, aligned in two NW-SE trending zones, probably controlled by the structure of the southern wall of Toya caldera (Figure 21) [Hernández et al., 2001; Miyabuchi et al., 2014, and references therein]. Usu has erupted every 30–50 years during the last three centuries with dome formation sometimes accompanying pyroclastic flows. Since the 1663 eruption, nine eruptions occurred, of which four were in the twentieth century, in 1910, 1943–1945, 1977–1982, and 2000. The composition of the erupted products evolved from rhyolitic to dacitic and the temperature of the magma apparently increased with time, possibly due to sequential tapping from a stratified magma chamber replenished with high-temperature magma [Tomiya and Takahashi, 2005; Aoyama et al., 2009; Matsumoto and Nakagawa, 2010]. Unlike the previous eruptions, the 1977–1982 eruption occurred on the summit crater: after only 36 h of precursory felt earthquakes, a large dacitic ash column of the subplinian type rose on the morning of 7 August 1977. More than 10 eruptions, including 4 large ones, were repeated during the first week. In this beginning stage numerous faults were found on the summit crater [Newhall and Dzurisin, 1988; Aoyama et al., 2009].

A new crisis at Usu started on 27 March 2000, with a progressive increase in the number and magnitude of the earthquakes. On 28 March, 599 volcanic earthquakes and 5 low-frequency events below Usu were followed, on 29 March, by 1629 volcanic earthquakes and 164 low-frequency events, with largest magnitudes of M = 3.7 and M = 4.1. On the afternoon of 29 March, GPS sensors also recorded the first signs of inflation. On 30 March, 2454 volcanic earthquakes and 326 low-frequency events were accompanied by a sharp increase in the deformation rate. On 31 March, 788 volcanic earthquakes and 139 low-frequency events occurred under a decreasing trend in the seismicity, with slow deformation near the volcano and fast deformation near the site of forthcoming eruption. On the early afternoon of 31 March, a VEI 2 eruption started on the NW slope of Mount Usu and seismic activity decreased [Jousset et al., 2003, and references therein; Zobin et al., 2005; Aoyama et al., 2009]. The preeruptive earthquakes, reaching a maximum magnitude of 4.6, clustered beneath the summit and southern slope of the volcano, mainly at depth between 4 and 7 km; the epicenters did not reach the sites of the forthcoming eruption but stopped at a distance of 1–3 km [Zobin et al., 2005]. The long-period tremors, located at a depth of 5 km and with amplitude variation correlated with the uplift rate, have been related to the flow-induced vibration of a magma chamber; this was located at ~5 km depth SW of the summit area and fed by a 10 km deep magma chamber [Yamamoto et al., 2002]. The shallower part of the precursory seismic activity, above the magma chamber at ~5 km of depth, reflects preeruptive magma movements and may be divided into three parts: (1) a subvertical distribution, indicating magma ascent to beneath the summit of the volcano, at a depth from 4 to 2 km; (2) a northward migration that indicates magma movement shallower than 2 km producing the eruption; and (3) a horizontal southward migration, possibly indicating intrusion of sill at 3 km of depth beneath the summit [Onizawa et al., 2007].

Magmatic CO₂ flux from the summit area increased from 120 t/d (September 1998) to 340 t/d (September 1999), and probably higher still before the 31 March 2000 eruption, and then decreased sharply to 39 t/d in June 2000, 3 months after the eruption. The change in CO₂ flux and seismic observations suggests that before the eruption advective processes controlled gas migration toward the surface [Hernández et al., 2001]. Water levels in wells decreased from October 1999, 6 months prior to the eruption, probably reflecting dilatation while magma was still deep. Then, abrupt and large rises in well levels were observed 3 days before the eruption, simultaneously with the increase of seismic activity. These water-level rises are caused by crustal compression as magma neared the surface. No variations in temperature and chemical composition of thermal waters and fumaroles were evident prior to the eruption [Shibata and Akita, 2001; Matsumoto et al., 2002; Shibata et al., 2008]. During the early phase of the eruption, uplift of the Usu edifice was geodetically detected. From 3 to 29 April, inflation was observed with vertical and horizontal displacements each exceeding 20 m. During the period of major activity from 3 to 5 April, uplift and N-S spreading velocities reached 3.3 and 2.7 m/d, respectively [Tobita et al., 2001]. Faulting and the main explosive eruptions did not take place in the highest uplifted area, but along the margin. This suggests that the faulting and explosive activities diverged from the main magma body responsible for the highest uplift. Such a significant faulting and deformation of the ground surface has been rarely known during volcanic eruptions [Yamagishi et al., 2004].

In synthesis, the 2000 Usu eruption had months of precursors in the water level and CO₂ gas flux, and then a very few days of immediate seismic and deformation precursors. Magma may have been added to the shallow chamber at ~5 km of depth since September 1998, when an increase in the CO₂ flux started, and then rose from the magma chamber to the surface in a time span as short as 5 days.

An overall similar structural configuration to Toya is found in Aira, a caldera within the Kagoshima graben, on Kyushu Island, SW Japan. Aira belongs to the Ryukyu Volcanic Arc, driven by the subduction of the Philippine Sea plate beneath the Eurasian plate along the Ryukyu trench, ~200 km to the east. The caldera has produced many ignimbrites: the youngest and probably largest eruption, at ~22 ka, produced >400 km³ of magma, forming the current outline of the caldera [Aramaki, 1984]. The possible vent area for this eruption is the Wakamiko depression, in the NE part of Aira caldera. Sakurajima is an andesitic stratovolcano formed in the last 13 ka on Aira's southern rim [Sassa, 1956; Dvorak and Dzurisin, 1997]. The Aira caldera has produced >17 plinian eruptions, including a few from Sakurajima. Since 1955, Sakurajima has erupted frequently from the Minamidake crater, producing VEI 1–3 eruptions [Ishibashi et al., 2008; Roulleau et al., 2013; Kobayashi et al., 2013]. Since measurements began in 1890, the Aira-Sakurajima system has shown alternating inflation and deflation: when the summit craters of Sakurajima were inactive (as between 1964 and 1975), uplift of Aira caldera was generally observed, and vice versa [Eto et al., 1997]. Recent seismicity has been focused in three main areas: (a) below Sakurajima, at depths of 0–4 km; (b) below Aira caldera, at depths 4–14 km; and (c) in the area to the SW of Sakurajima, at depths of 6–9 km [Hidayati et al., 2007].

GPS and InSAR data from 1995 to 2007 indicate an overall uplift of Aira, with a magmatic source located at the center of the caldera, between 9 and 12 km depth [Rémy et al., 2007; Iguchi et al., 2008]. The GPS network also suggests a second source of deformation below the central crater of Sakurajima, at depth of 4–6 km [Iguchi et al., 2008, 2013]. Seismic activity, mainly consisting of VT events, gradually increased since 2002, focusing along a NE-SW trend from NE of Aira to SW of Sakurajima: one of the largest earthquake swarms, SW of Sakurajima, occurred between November 2003 and February 2004. The GPS baseline indicated coeval extension (report JMA, http://www.jma.go.jp/jma, MVAR March 2002). Eruptive activity at the Showa crater of Sakurajima resumed in June 2006, after 58 years of dormancy; eruptive activity at Showa from June 2010 to July 2011 was characterized by seismicity and repeated downward and upward tilts of the sides of Sakurajima (report JMA, http://www.jma.go.jp/jma, MVAR September 2014). This behavior is confirmed by InSAR data of the summit of Sakurajima, showing yearlong periods of subsidence alternating with uplift phases. InSAR data from June 2006 to April 2011 on Aira also show a general uplift of ~0.5 cm/yr, mostly on the edge of the central western portion of the caldera, from the northern part of Kagoshima Bay to the north flank of Sakurajima Volcano. This uplift is in agreement with other geodetic data [Rémy et al., 2007; Iguchi et al., 2013]. The inversion of the InSAR data suggests a deep source of magmatic origin (10 ± 5 km depth) that increases in volume of ~6 × 10⁶ m³/yr, located on the top of a large low-velocity area below Aira detected through regional seismic tomography.

In synthesis, the Aira-Sakurajima system provides a more active and restless version of the similar Toya-Usu system, also characterized by an active stratovolcano on the edge of a caldera; the larger activity at Aira-Sakurajima is testified by an open conduit at the continuously erupting Sakurajima and widespread deformation of Aira.

3.2.2.6 Krakatau and Tengger

The Krakatau complex lies along the Indonesian volcanic arc, in the Sunda Strait between Sumatra and Java. Krakatau appears to cycle through basaltic, basic andesitic, acidic andesitic, and dacitic eruptive phases, each cycle culminating in a massive destructive dacitic eruption [Camus et al., 1987]. Krakatau currently consists of four islands: Sertung, Panjang, Rakata, and Anak Krakatau. The first three are the remnants of the eruption and caldera collapse of the 1883 eruption, which resulted in >36,000 deaths, primarily due to the massive tsunamis inundating the surrounding coastlines; Anak Krakatau is a volcano island in the center of the Krakatau complex [Agustan et al., 2012, and references therein]. Petrological and geophysical data suggest the existence of at least two main magma reservoirs at depths of 3–7 km and 7–12 km beneath Anak Krakatau [Dahren et al., 2011]. Anak Krakatau has been the site of frequent mafic eruptions, all with VEI ≤ 2, since 1927.

Considering the last decades of activity, Krakatau produced 11 eruptions: 7 eruptions from 1992 to 2001, with an average interval of 10 months and 4 eruptions from 2007 to 2012, with an average interval of 7 months. In general, seismicity, recorded several months before each eruption, was the most common precursor. For example, the eruption begun on 7 November 1992 was preceded by an increase in seismicity for a week [Agustan et al., 2012]. Deformation, detected through InSAR, showed significant changes before and after the 2007 eruption, following the onset of seismic activity. This eruptive event, characterized by Strombolian activity with ash columns 1 km high, as well as pyroclastic and lava flows, lasted from the end of October 2007 to August 2008. The eruption was preceded by an increase in seismic activity from June 2006 and, from 3 months before its onset, by a complex pattern of ground deformation: inflation up to 4 cm, together with subsidence around the crater, were detected [Agustan et al., 2012]. An eruption, with the same preeruptive seismic activity and surface deformation, took place between October 2010 and September 2012 [GVP report, 12/2012 (BGVN 37:12)].

Preeruptive behavior similar to that of Anak Krakatau, characterized by minor seismicity and surface deformation in an open conduit caldera, has been observed at Bromo Volcano, in the Tengger caldera in east Java. Tengger is a 16 km wide caldera, reaching 2300 m asl, the youngest among the five belonging to a volcanic complex active in the last 820 ka. Bromo recently erupted in 1995 (VEI 1), 2000 (VEI 2), 2004 (VEI 2), and 2010 (VEI 3), with a similar preeruptive behavior, where seismicity and deformation are short-term precursors, active only just prior to the eruption. The usual deformation of the volcano before (inflation) and after (deflation) any eruption is relatively small, i.e., typically in order of a very few centimeters [Abidin et al., 2004]. Preeruptive seismicity is commonly detected, with tremor signals largely due to flow of gases and steam through an irregularly shaped conduit and shock signals attributed to fissuring of the conduit wall due to the high pressure inside [Gottschammer and Surono, 2000; GVP report, 07/2004 (BGVN 29:07)].

4 Discussion

4.1 A Synthesis of the Database: Identifying the Most Recurrent Types

All of the 42 monitored calderas in the last decades have shown at least one period of unrest, based on variations in any of the monitoring parameters (Figure 3). This indicates that unrest is a normal condition which should be experienced by any caldera on a time span of a very few decades. Overall, 110 of 166 unrest episodes were eruptive, and most (72 cases) occurred at mafic calderas. Conversely, noneruptive unrest episodes are more uniformly distributed between mafic and felsic calderas (Table 1). This suggests a tendency for mafic calderas to erupt more easily during unrest; however, it must be noted that the number of mafic calderas considered in our analysis is double that of felsic ones.

The described cases suggest several types of unrest (Tables 2 and 3). A first general distinction may be made based on the overall composition, distinguishing the behavior of mafic and felsic calderas.

Table 2. Summary of the Unrest Types at Calderas From 1988 to 2014

Table 3. Average Duration of Various Types of Unrest During the Past 126 Years (Sample Size in Parentheses; Updated From Newhall and Dzurisin [1988])

Type of Unrest	Magma of Caldera-Forming Eruption
Type of Unrest	Mafic	Felsic	Unknown
Local earthquakes	Days/weeks (193)	Weeks to months (230)	Days to weeks (15)
Earthquake swarms	Days/weeks (193)	Weeks to months (230)	Days to weeks (15)
Tremor	Hours to days (79)	Days to weeks (43)	−(2)
Uplift	Months to years (72)	Months to years (69)	Years to decades (13)
Subsidence	Weeks (27)	Months to years (36)	−(0)
Tilt	Weeks to months (39)	Months (14)	−(0)
Horizontal strain	Months (7)	Months to years (5)	−(0)
Ground fracturing	Days to weeks (8)	Weeks (20)	−(1)
Magnetic changes	−(3)	−(2)	−(0)
Gravity changes	−(12)	−(9)	−(0)
Thermal changes	Months (4)	Months (40)	Months (6)
Fumarole, gas changes, spring, lake changes	Months (42)	Weeks to months (108)	Months (16)

Mafic calderas (including Galapagos, Hawaii, Izu-Oshima, Aoba, Karthala, Masaya, Poas, and Pacaya) show cycles of inflation, last-minute seismicity, eruption, deflation, suggesting a consistent behavior. Seismicity is usually observed on the short-term, immediately before the eruptions, probably due to the common opening of the conduit and the shallowness of the magma. Deviations from this quite predictable behavior are observed at Sierra Negra (showing longer and composite unrest, with larger displacement and erupted volumes), Dolomieu and Miyakejima (during the lateral intrusion of magma, which may generate caldera collapse above the withdrawn reservoir) and Mauna Loa and Taal (both noneruptive during the period of our study). Other mafic calderas, as Askja or Krafla, have been experiencing long-term deflation without any eruption.

Unrest at felsic calderas tends to be longer (decades or even longer) but increase sharply hours to days before eruption. The final increase in unrest is, in at least some cases, triggered by sudden injection of fresh magma into the system. For some calderas, such as Santorini, short magma accumulation after long repose times has been inferred to occur at various timescales and proportionally to the volume of the accumulated magma so that longer repose times are accompanied by relatively longer and larger replenishments. Large felsic calderas (>10 km wide) and without a central edifice often exhibit multiple episodes of unrest, suggesting multiple magma intrusions into, through, or near the felsic magma reservoir. Some of these intrusions have erupted within the past century (Aso, Aira-Sakurajima, Toya-Usu, Krakatau, Tengger-Bromo, and Rabaul), whereas others have not (Campi Flegrei, Long Valley, Yellowstone, Colli Albani, and Iwo-Jima).

Important interaction between the magmatic system and a hydrothermal source has been highlighted at several felsic calderas, including Campi Flegrei, Yellowstone, Pinatubo, and Aso [e.g., Unglert et al., 2011; Chiodini et al., 2012a]: in most cases, fluid transfer occurs vertically from the deeper magmatic source to the shallower hydrothermal system; however, in the case of Yellowstone, the transfer of fluids occurs mainly laterally, to and from the main hydrothermal area outside the caldera [Chang et al., 2007]. The interaction between the magmatic and hydrothermal system introduces important variations in the latter, usually amplifying the changes in the monitoring parameters at the surface.

Interactions between regional tectonic earthquakes and the onset, or intensification, of unrest have also been highlighted at several felsic calderas, as at Long Valley, Campi Flegrei, and Rabaul [Linde et al., 1994; Johnson et al., 2010c; Lupi et al., 2015]. However, at Rabaul the response of the caldera system to regional earthquakes has been nonlinear, with more distant earthquakes affecting unrest rather than nearer earthquakes with similar magnitude.

In some cases, the size of the felsic magmatic system may play an important role in determining the behavior of the caldera during unrest. Smaller magmatic systems are less effective at buffering magmatic intrusions. For example, at Pinatubo the intrusion of fresh mafic magma was buffered for a few months by the residual dacite, but not for years, as is often the case at larger calderas (Yellowstone and Long Valley). Larger felsic magmatic systems are also associated with well-developed and active hydrothermal systems, generating more complex unrest patterns, as at Yellowstone. In some cases, active stratovolcanoes may also develop on the edge of these large felsic calderas, as at Toya-Usu and the more active and restless Aira-Sakurajima.

The rates of the indicators of unrest may vary significantly. Lower rates (uplift of mm/yr, sporadic seismicity) are observed at Colli Albani [e.g., Riguzzi et al., 2012]; higher rates (uplift of centimeters to meters per year, substantial degassing, continuous seismicity) are observed at Campi Flegrei, Long Valley, Yellowstone, and Iwo-Jima [Hill et al., 2003; Ukawa et al., 2006; Chang et al., 2007; Del Gaudio et al., 2010], or at the Laguna del Maule caldera, southern Andes [Feigl et al., 2014; Singer et al., 2014]. Felsic calderas are more prone to high rates of unrest, especially where there is a hydrothermal system and plugged or semiplugged conduit(s). However, except for unrest at Sierra Negra [Chadwick et al., 2006], higher rates are seldom observed at mafic calderas.

The unrest types summarized in Table 2 suggest that while mafic calderas are usually characterized by repeated cycles of inflation-eruption-deflation, felsic calderas show more complicated behaviors, also involving the activity of the hydrothermal system, with isolated or more frequent injections of magma, only in part erupting. It may be thus postulated that under similar replenishment rates and frequencies, felsic calderas release magma and erupt less frequently than mafic calderas. Different factors (higher depth and larger size of the magma chamber, low density and high viscosity, and presence of a hydrothermal system) all contribute to this filter or buffer effect, which retards magma from erupting from felsic calderas.

When available, the measured variations in the geophysical, geodetic, and degassing parameters are usually coupled so that increasing rates of an indicator correspond to increasing rates of other indicators (Table 2); increases in seismicity, accompanied by contemporaneous uplift, are the most commonly detected variations. Minor deviations from this pattern have been observed in deflating volcanoes with minor preeruptive seismicity (Aso), in the presence of volcanoes with open magmatic conduit, where the seismicity may be poor or absent and, at times, at Taal and Yellowstone. The partly noncoupled behavior of Taal and Yellowstone may be explained by interactions between the magmatic and the hydrothermal system [Bartel et al., 2003; Chang et al., 2007]. Except for these minor variations, most unrest is characterized by an overall coupling of the increasing monitoring parameters, i.e., by a linear behavior of the indicators of the system during unrest. There are a few cases of nonlinear behavior between the variation of the deformation of the caldera and the occurrence of an eruption. These include Aso (erupting while deflating) or Okmok (erupting during a stasis in the deformation) [Lu and Dzurisin, 2014]. Thus, deformation may not always be a reliable indicator to forecast any impending eruption. A further limitation in interpreting the deformation pattern to forecast eruptions regards the discrepancy between the location of the maximum uplifted area and that of the eruptive vent, as at Usu in 2000 and Sierra Negra in 2005 [Yamagishi et al., 2004; Jonsson et al., 2005].

While all the caldera eruptions we considered are preceded by unrest, not all the unrest episodes culminate in eruptions. Therefore, our data set suggests that unrest is a necessary but not sufficient condition to have an eruption [e.g., Newhall and Dzurisin, 1988; Moran et al., 2011]. In particular, ~1/3 of cases of unrest from our data set, equally at felsic and mafic calderas, were not followed by any eruption. In some cases, isolating and considering as independent each unrest episode within long-term restless behavior may be misleading. This is the case of Rabaul, when, in retrospective, the strong unrest in the early 1980s primed the system and the lesser unrest in the early 1990s set off the eruption [Johnson et al., 2010c]. Especially at felsic restless calderas with plugged or semiplugged conduits, repeated intrusions may accumulate and increase the hazard. We should look at the history and development of multiphase unrest, rather than focusing on single episodes. Rabaul also provides a further complexity, as the 1994 eruption occurred simultaneously from two different vents along the caldera inner ring fault [Johnson et al., 2010c].

4.2 A Comparison to Newhall and Dzurisin [1988]

Two of the major conclusions of an earlier review of caldera unrest [Newhall and Dzurisin, 1988] are reaffirmed by the present study. One, that most large calderas had experienced unrest within historical time is reaffirmed by our finding that in the past 25 years and at 42 calderas for which good monitoring data were available, all 42 experienced unrest in 1 year or another and often for multiple years. Another, that unrest at calderas had multiple causes, is reaffirmed by our finding that magma intrusions, hydrothermal systems, and regional tectonics all contributed to caldera unrest of the past 25 years.

The present study goes a step further in causal attribution that all episodes of unrest for which cause could be established involved magma intrusion. Hydrothermal systems may also have been engaged, and regional tectonics may also have contributed, but magma was the fundamental driver of unrest. Partly, this conclusion reflects our selection of only those calderas that had good monitoring, and partly, it reflects newer, better understanding of the origins of phenomena like elevated CO₂ outgassing, deep long-period earthquakes, and others.

Both studies, Newhall and Dzurisin [1988] and ours, suggest that large caldera systems, especially silicic ones, have a high capacity to buffer incoming disturbances (magma resupply and tectonic disturbances) without progressing to eruptions. However, a higher percentage of our case studies progressed to eruption than reported in Newhall and Dzurisin. We attribute this to more emphasis and more detailed counting of unrest and eruptions at mafic calderas in this paper, compared to that in Newhall and Dzurisin [1988].

A new insight in the current work pertains to the important role of differences in magma degassing from one caldera to another, especially felsic ones. In reviews that emphasized arc stratovolcanoes, Newhall [2004], Newhall [2015], and Whelley et al. [2015] distinguished between volcanoes that release most of their gas passively, between eruptions (variously termed open conduit, open vent, or leaky) versus those that trap incoming gas and accumulate it in the magma chamber even in excess of saturation (variously termed closed conduit, closed vent, tight, or plugged). In the case of plugged volcanoes like Pinatubo, accumulated gases are released in large volume during explosive eruptions. Intermediate behavior might be called semiplugged [Whelley et al., 2015].

How might different time histories of degassing affect caldera unrest? Most mafic calderas have continuous degassing and relatively frequent, small and/or effusive eruptions. In the spectrum of open versus plugged, most of these are open systems, and gas never accumulates at magma storage depths in excess of saturation. The more open a system, the more subtle its geophysical indicators, because little fresh fracturing occurs, and only modest magma pressures develop. At felsic calderas, though, gas budgets are more difficult to judge. Most of the eruptions that have occurred at these calderas during the past 25 years have also been small (VEI < 4) and have occurred from mafic to intermediate-composition cones within or on the rims of these calderas (e.g., Anak Krakatau within the Krakatau caldera; Nakadake within Aso Caldera, Sakurajima on the rim of Aira Caldera). Most of these subsidiary cones show open-conduit behavior, but how about the calderas as a whole? The earlier studies of Newhall and others relied on measured intereruption and syneruption degassing of SO₂, but at large calderas as a whole (i.e., across the entire area of a caldera), intereruption SO₂ may be seriously underestimated if it is absorbed (“scrubbed”) into the groundwater system of the caldera. Syneruption SO₂ may also be underestimated because large-scale eruptions have not occurred since 1979, the period of global SO₂ monitoring by satellites. As a result, when we try to classify large felsic calderas in the open versus plugged spectrum, uncertainties are high. For example, are Campi Flegrei and Yellowstone open, semiplugged, or nearly plugged? We know that in both cases there is substantial CO₂ outgassing, but what percentage does this represent of the CO₂ being supplied to the system at depth? Ongoing studies on Campi Flegrei suggest that a significant part of the magmatic gases supplied to the hydrothermal system are released at the surface through major fault systems, with a dramatic increase in the last decade (Figure 22) [Acocella, 2010; G. Chiodini, personal communication, 2015]. Perhaps the best we can do is to assume that large felsic calderas with frequently erupting, openly degassing subsidiary vents are effectively releasing all incoming volatiles, while those without such surface manifestations may be accumulating volatiles at magma storage depths unless we can measure high magmatic CO₂ fluxes, as at Campi Flegrei. Finally, one caldera in our study (Iwo-Jima) seems to be an end-member, leaking most or all incoming volatiles even without frequent small explosive eruptions. Seismicity is relatively minor, but inflation is dramatic, reflecting shallow intrusions that remain as intrusions when their gas bleeds off.

An important implication of the preceding discussion is that geophysical and geochemical indicators (precursors) of eruptions at calderas will not always show a slow, progressive buildup to eruption. Often, the magma and hydrothermal systems of calderas (especially felsic calderas) will buffer small intrusions, and unrest will stop without eruption. Yet the same magma intrusions might remain molten and push the system closer to critical state, from which a final push toward eruption might occur in just hours or days (e.g., Campi Flegrei 1538, Rabaul 1994, and Okmok 1997 and 2008). Another implication—especially at continuously degassing calderas—is that the continuous bleed of gas limits the explosive potential of the caldera. New, large, caldera-scale eruptions are unlikely to occur where gas bleeds off continuously—regardless of the repose time since the last large eruption. In contrast, calderas that do not show much degassing might be accumulating gas and explosive potential in proportion to the repose since the previous large eruption.

4.3 Trying to Understand Processes

The calderas in our study have shown a large variability in the manifestation of unrest, including some specific and unexpected behaviors. However, despite any variability, some very general and basic features of the considered unrest types may be also recognized. The above discussions suggest that the unrest type at a given caldera depends strongly upon its composition, here simply schematized as mafic or felsic, and the degree to which incoming magma degasses freely, here schematized as open, semiplugged, or plugged.

Section 4.1 has shown that mafic calderas show more regular behavior with inflation, last-minute seismicity, eruption, and deflation, and felsic calderas show more complex and dramatic variations in the indicators and many episodes of unrest that do not culminate in eruption. These different behaviors probably reflect lower magma viscosities, fewer density barriers to magma ascent, and generally smaller size of mafic systems.

A caldera with plugged conduit does not allow gas or magma to rise easily toward the surface; this has been, for example, the condition of Pinatubo Volcano before the 1991 eruption; a caldera with semiplugged magmatic conduits releases gas to the surface but not magma; this is, for example, the current state of Campi Flegrei, characterized by significant CO₂ emissions; fully open conduits are only found in mafic systems, where both gas and magma can rise freely to the surface, as at Erta Ale caldera, Afar, or Kilauea (Figure 22).

The effects of both the reservoir magma composition and the plugging of conduits seem to explain most behavior of calderas during unrest in the last decades (Figure 23). Except for Sierra Negra (a mafic caldera with “felsic” behavior), it appears that these parameters may adequately describe the main types of caldera unrest. A first distinction, based on the mafic versus felsic composition, has been already made in section 4.1. The degree of plugging or openness of the system, discussed in section 4.2, adds further information to this distinction, with the specification that while mafic calderas may easily reach the open end-member, conduits at felsic calderas are generally plugged or semiplugged.

In mafic calderas, unrest appears, in general, to be subtler and less pronounced, especially with an open system, which usually results in low seismicity, as, for example, at the western Galapagos calderas, Dolomieu, or Kilauea. These open calderas may also erupt more frequently, even though with limited magnitude. Examples of closed mafic calderas, without recent eruptive activity, may be found in Iceland (Katla). In between are the semiplugged calderas, as Mauna Loa or Taal.

In the case of felsic calderas, well-plugged systems are usually smaller, with isolated and short unrest episodes, erupting infrequently, as Pinatubo, or Changbaishan. Semiplugged systems are usually larger, restless over decades or centuries, and erupt even less frequently, as Yellowstone, Long Valley, or Campi Flegrei. At these semiplugged calderas, the common association between multiple unrest and resurgence suggests repeated intrusions without eruptions, explaining observed resurgences. The thermal weakening and widespread fracturing of the crust overlying these larger magma chambers may enhance the ascent and leakage of the magmatic gases, repeatedly emplacing gas-poor intrusions at depth and thus providing the basic conditions for resurgence. At the same time, high viscosity and low density of silicic magma chambers impedes or buffers access of fresh mafic magma to the surface. Rabaul exhibits behavior intermediate between plugged Pinatubo and the unusually open felsic system of Iwo-Jima.

In summary, composition of reservoir magma and whether or not magma in conduits rising from the reservoir solidifies between eruptions (as reflected by degassing patterns) offer a general key to interpret and forecast the outcome of unrest. In Figure 24, four end-members or types of unrest may be defined. These diagrams are a conceptual representation of different magmatic processes below calderas, particularly those relative to the magma heat loss and crystallization at depth, as well as the overall extrusive to intrusive ratio of the system and its consequent capability to trigger a specific type of unrest. In general, mafic calderas with open conduits provide the easiest conditions for the continuous, aseismic and moderate release of magma, anticipated by minor, even though repeated, unrest. At the opposite are the felsic calderas, with conduits both plugged and semiplugged. Where conduits are plugged, most of the energy is stored within the magmatic system and violently released through infrequent eruptions, anticipated by seismically active unrest. Semiplugged systems, including some of the largest magmatic systems on Earth (Yellowstone and Long Valley) are sites of significant volatiles and magma production: however, while the volatiles are lost through degassing, most of the magma remains stored in repeated intrusions (schematically reported as sills in Figure 24), with a restless behavior testified by uplift and seismicity and, on the longer term, by resurgence.

Mafic calderas are likely to have minor and rapid precursory phenomena (especially seismicity) related to the direct ascent of mafic magmas. Felsic calderas are more likely to have repeated noneruptive unrest that represents a recharging or underplating of a shallow reservoir. When a felsic caldera is fully primed, though, unrest leading to an eruption may be quite fast (as at Rabaul), as it represents the final and rapid ascent of gas-rich magma from a shallow overpressured reservoir that breaks open. This is consistent with a preliminary statistical analysis of the caldera unrest database, which suggests that eruptive unrest is much shorter (usually <10 months) than noneruptive unrest [Di Lorenzo et al., 2015].

Seismicity appears the most effective indicator to forecast both the possible location of the eruptive vent and the onset of the eruption, especially during the final runaway phase, when earthquakes cluster around the propagating feeder dike. Surface deformation is usually reliable, even though at times misleading in forecasting both the location (as at Sierra Negra) and the onset of eruption (deflating or stationary rates at Aso and Okmok). Degassing is usually limited to the most fractured areas, typically with uneven distribution within the caldera, and may be influenced by interactions with the water table.

4.4 Possible Future Advances

Understanding of the processes behind caldera unrest may significantly benefit from several advances, including the following:

Better understanding of the overall volcanic history of each caldera, identifying both major caldera forming and intracaldera activity. This includes the eruptive frequency, the location of the vents, and the inferred VEI.
Better understanding of the structural framework of the caldera and its control on magma ascent. Calderas, like all other volcanoes, are leaking valves on the Earth surface and as such are important to understand not only the flux of magma but also the external controls and regulations. Several studies have been recently highlighting the relationships between external regional earthquakes and caldera unrest [e.g., Takada and Fukushima, 2013; Lupi et al., 2015]. This sector of research definitely needs to be further explored.
Better understanding of gas budgets at calderas, to judge whether gas is accumulating at magma storage depths (e.g., as occurred at Pinatubo) or, alternatively, being bled off just as fast as it is supplied from depth. In turn, this requires estimation of volatile concentrations in resupplied magma, and rates at which magma is resupplied over annual, decadal, and longer time frames. Minimum volumes of magma resupply can be estimated from geodetic methods, from measurements of degassing, and from eruptive volumes. In addition, periodic surveys of CO₂ flux from whole calderas will be helpful, as would be comparisons of CO₂ flux from whole calderas and from their subsidiary cones.
Better understanding of shallow magma emplacement processes, as well as the associated monitoring signals which may be detected at the surface. Recent studies provide important insights on the factors controlling shallow magma propagation and emplacement in the crust, as well as the possible outcome [Geshi et al., 2010; Gudmundsson, 2011; Chanceaux and Menand, 2014; Macedonio et al., 2014; Kavanagh et al., 2015; Rivalta et al., 2015]. In future studies, attention should also be given at providing a possible hierarchy of the factors controlling magma propagation and emplacement as a function of the most likely boundary conditions at a given caldera.
Better definition of the role of any hydrothermal system during unrest. In particular, it should be defined whether and under which conditions a given hydrothermal system may amplify, buffer, or have negligible impact on the magma reservoir dynamics. On the one side, the limited systematic monitoring of composition and flow rate of fumarole and thermal water at calderas encourages the deployment of automatic stations able to record chemical and isotopic variations of hydrothermal fluids and their fluxes. On the other side, advanced numerical modeling, taking into account for the coupled dynamics of magma reservoirs and hydrothermal systems, as well as for the supercritical conditions of fluids within the latter, is fundamental to better detect and understand unrest, especially at felsic calderas.
Better definition of the patterns of the monitoring indicators and their possible relationships during unrest, with particular attention at identifying any general behavior.
Better link of the short-term processes (years or decades), operating during unrest and detected through monitoring, with the longer-term ones (hundreds to thousands of years), observed through the geological record, most notably resurgence. Resurgence is one of the least understood processes in volcanology: while there is a general and shared consensus on how calderas form, there is no similar broadly accepted agreement on how a caldera floor may uplift up to a few kilometers. Recent studies highlight episodic uplift during unrest, even with remarkable rates, as at Iwo-Jima, Tanna, Campi Flegrei, Ischia, and Toba [Ukawa et al., 2006; Ozawa et al., 2007; Vezzoli et al., 2009; Del Gaudio et al., 2010; Brothelande, 2015; De Silva et al., 2015]. If resurgence results from multiple intrusions without eruptions, what are the conditions required to inhibit magma from erupting, and leave it stored at depth, uplifting a volcano? One important factor may be degassing, but more data on degassing are needed to test this proposition.

5 Conclusions

Our study suggests the following points.

Where established, the root cause for unrest is always magmatic; while external earthquakes or hydrothermal processes may definitely enhance unrest, no unrest was of purely hydrothermal or tectonic origin; magma always appears as the crucial ingredient, though in most cases, magma may also induce changes in the hydrothermal system, supplying this with fluids and energy.
An interpretive classification of unrest invokes two spectra—compositional (mafic to felsic) and the state of magma conduits feeding from the magma reservoir(s) to the Earth's surface (from fully plugged, through semiplugged, to open); in particular, magma and gas in open conduits can rise and erupt freely; magma in semiplugged conduits is more viscous and erupts less frequently yet still allows some gas to rise and escape; plugged conduits allow neither magma nor gas to escape.
In mafic calderas, unrest appears to be subtler and less pronounced, especially with an open magmatic system, which usually results in low seismicity. These open calderas may also erupt more frequently, even though with limited magnitude. Therefore, mafic calderas with open conduits provide the easiest conditions for the continuous, aseismic and moderate release of magma, anticipated by minor, even though repeated, unrest.
In felsic calderas, well-plugged systems are usually smaller, with isolated and short unrest episodes, erupting infrequently. Here most of the energy is stored within the magmatic system and violently released through infrequent eruptions, anticipated by seismically active unrest. Semiplugged felsic calderas include some of the largest magmatic systems on Earth, sites of significant volatiles and magma production. These are restless over decades or centuries and erupt even less frequently, showing unrest with uplift, seismicity and degassing, and, on the longer-term, resurgence. The common association between multiple unrest and resurgence suggests repeated intrusions without eruptions. The thermal weakening and widespread fracturing of the crust overlying these larger magma chambers may enhance the ascent and leakage of the magmatic gases, repeatedly emplacing gas-poor intrusions at depth and thus triggering resurgence. At the same time, high viscosity and low density of silicic magma chambers impedes or buffers access of fresh mafic magma to the surface.

Glossary

Caldera: A wide topographic depression, usually much larger than the diameter of the conduits that feed eruptions, formed during or soon after an eruption by subvertical subsidence into a partly drained magma reservoir.
Dike: Subvertical sheet-like intrusion transferring magma from a reservoir to the surface.
Felsic caldera: A caldera erupting dominantly silicic (trachytic, andesitic, dacitic, and rhyolitic) or dominantly silicic evolved products. Felsic calderas may be several tens of kilometers wide, with collapse of a few kilometers downward and may exhibit one or more resurgences.
Hydrothermal system: Heated magmatic and nonmagmatic fluids, usually in the liquid phase, within fractured and porous rocks, usually located between the magma chamber and the surface.
Ignimbrite eruption: Explosive eruption producing pumiceous, ash-rich deposits of pyroclastic density currents.
Immediately preeruptive (or runaway) unrest: Any distinctive, diagnostic unrest, often a rapid acceleration, just a few days to hours before the onset of an eruption. This unrest is the expression of the propagation and integration of the feeder dike from a magma reservoir to the surface.
Mafic caldera: A caldera with dominantly basaltic or primitive products. Mafic calderas are usually of limited extent (a few kilometers wide) and subsidence (hundreds of meters), with rare resurgence.
Magma reservoir: Portion of the Earth's crust where magma accumulates and resides temporarily on its way from the upper mantle to the surface.
Monitoring system: Permanent and/or mobile instrumentation network devoted at detecting and tracking geophysical or geochemical activity (mainly seismicity, ground deformation, microgravity, chemical and isotopic composition of fumaroles, and degassing of fumaroles and hydrothermal sites).
Open conduit: Where magma in the conduit(s) of the previous eruption allows free ascent of magma and gas. Often, magma is convecting in the open conduit and, near the surface, becomes a permeable foam that releases most or all of the volatiles supplied to the reservoir between eruptions. Small magmatic eruptions may be frequent.
Phreatic eruption: Explosive eruption due to the heating and pressurization of ground or surface water by magma, without direct contact.
Plugged conduit: A conduit where magma of the previous eruption solidifies before the next eruption, preventing newly supplied gas and magma from escaping. Some conduits are plugged between some eruptions and open between others (e.g., at Vesuvius), depending on magma viscosity and the rates of magma resupply.
Restless caldera: A caldera which has undergone a prolonged (several years to centuries) variation in any monitoring parameter.
Resurgence: Evident uplift (≥10¹ m) of part of a caldera floor and fill, commonly forming a central dome with an apical graben. Much more common in felsic calderas than in mafic calderas.
Ring-fault system: An annular, elliptical or polygonal system of faults around a caldera enclosing the central subsided area.
Ring dike: Vertical to steeply outward dipping circumferential pyroclastic or magmatic intrusion commonly along a caldera ring-fault system.
Semiplugged conduit: Where magma in the conduit is viscous and only occasionally able to erupt but has fractures or slow convection that allow significant degassing into a hydrothermal system or into the atmosphere.
Sill: Subhorizontal sheet-like intrusion allowing the emplacement and accumulation of magma, also at very shallow crustal levels.
Unrest: Deviation of seismicity, deformation, gas emission, and/or other geophysical and geochemical activities from normal baseline(s) to elevated activity. Probabilities of eruption increase during unrest.
VEI or volcanic explosivity index: Parameter commonly used in volcanology to estimate the magnitude of an eruption, with 0 < VEI < 8 [Newhall and Self, 1982].

Acknowledgments

Editor F. Florindo provided effective and prompt support in handling the manuscript. Reviewers M. Moldwin, J. Cole, and G. Chiodini provided helpful suggestions, improving the study, and are gratefully acknowledged. N. Geshi and J. Ruch provided useful suggestions and support. S.H. Potter and L. Sandri provided helpful discussions. E. Fujita and M. Ukawa kindly provided material for a figure. Funded by DPC-INGV V2 project (V.A. responsible). Any user can access the data of this work (included in the supporting information, with the related explanatory notes) by contacting the corresponding author V. Acocella ([email protected]).

The Editor on this paper was Fabio Florindo. He thanks Jim Cole, Giovanni Chiodini, and one anonymous reviewer.

Supporting Information

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Citing Literature

Volume53, Issue3

September 2015

Pages 896-955

Filename	Description
rog20079-sup-0001-supplementary.docxWord 2007 document , 18.8 KB	Text S1
rog20079-sup-0002-supplementary.xlsxapplication/unknown, 98 KB	Database S1

An overview of recent (1988 to 2014) caldera unrest: Knowledge and perspectives

Abstract

Key Points

1 Introduction

2 The Database

3 Types of Caldera Unrest

3.1 Unrest Without Eruptions

3.1.1 Mafic Calderas: Representative Types

3.1.1.1 Askja and Krafla

3.1.1.2 Mauna Loa

3.1.1.3 Taal

3.1.2 Felsic Calderas: Representative Behaviors

3.1.2.1 Campi Flegrei

3.1.2.2 Long Valley

3.1.2.3 Yellowstone

3.1.2.4 Colli Albani

3.1.2.5 Santorini and Nisyros

3.1.2.6 Changbaishan

3.2 Unrest With Eruptions

3.2.1 Mafic Calderas: Representative Types

3.2.1.1 Fernandina, Cerro Azul, and Sierra Negra

3.2.1.2 Kilauea

3.2.1.3 Okmok

3.2.1.4 Recent Collapsing Calderas: Miyakejima and Dolomieu (Piton de la Fournaise)

3.2.1.5 Other Mafic Calderas

3.2.2 Felsic Calderas: Representative Types

3.2.2.1 Rabaul

3.2.2.2 Pinatubo

3.2.2.3 Iwo-Jima

3.2.2.4 Aso

3.2.2.5 Toya (Mount Usu) and Aira (Sakurajima)

3.2.2.6 Krakatau and Tengger

4 Discussion

4.1 A Synthesis of the Database: Identifying the Most Recurrent Types

4.2 A Comparison to Newhall and Dzurisin [1988]

4.3 Trying to Understand Processes

4.4 Possible Future Advances

5 Conclusions

Glossary

Acknowledgments

Supporting Information

References

Citing Literature

Figures

References

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