Persistent activity and violent strombolian eruptions at Vesuvius between 1631 and 1944

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Journal of Volcanology and Geothermal Research 170 (2008) 167 – 180 www.elsevier.com/locate/jvolgeores

Research paper

Persistent activity and violent strombolian eruptions at Vesuvius between 1631 and 1944 Roberto Scandone ⁎, Lisetta Giacomelli, Francesca Fattori Speranza Dipartimento di Fisica, Università Roma Tre, Via Vasca Navale 84, 00146, Roma, Italy Received 4 June 2007; accepted 30 September 2007 Available online 24 October 2007

Abstract During the period 1631–1944, Vesuvius was in persistent activity with alternating mild strombolian explosions, quiet effusive eruptions, and violent strombolian eruptions. The major difference between the predominant style of activity and the violent strombolian stages is the effusion rate. The lava effusion rate during major eruptions was in the range 20–100 m3/s, higher than during mild activity and quiet effusion (0.1–1 m3/s). The products erupted during the mild activity and major paroxysms have different degree of crystallization. Highly porphyritic lava flows are slowly erupted during years-long period of mild activity. This activity is fed by a magma accumulating at shallow depth within the volcanic edifice. Conversely, during the major paroxysms, a fast lava flow precedes the eruption of a volatile-rich, crystal-poor magma. We show that the more energetic eruptions are fed by episodic, multiple arrival of discrete batches of magma rising faster and not degassing during the ascent. The rapidly ascending magma pushes up the liquid residing in the shallow reservoir and eventually reaches the surface with its full complement of volatiles, producing kilometer-high lava fountains. Rapid drainage of the shallow reservoir occasionally caused small caldera collapses. The major eruptions act to unplug the upper part of the feeding system, erupting the cooling and crystallizing magma. This pattern of activity lasted for 313 y, but with a progressive decrease in the number of more energetic eruptions. As a consequence, a cooling plug blocked the volcano until it eventually prevented the eruption of new magma. The yearly probability of having at least one violent strombolian eruption has decreased from 0.12 to 0.10 from 1944 to 2007, but episodic seismic crises since 1979 may be indicative of new episodic intrusions of magma batches. © 2007 Elsevier B.V. All rights reserved. Keywords: Vesuvius; permanent activity; violent strombolian eruption

1. Introduction Mount Vesuvius is located in Southern Italy near the city of Naples; its products all postdate the 39 ka Campanian Ignimbrite (Scandone et al., 1991; Brocchini et al., 2001). Vesuvius has grown mostly on volcanic, alluvial and marine sediments that filled up the graben formed during the Pliocene and Pleistocene by the subsidence of Mesozoic carbonate platforms that make up the basement of the Campanian plain; this platform now lies about 2 km below the volcano (Ippolito et al., 1973; Scandone et al., 1991). The volcano is formed by an older volcanic edifice: (the “Somma”) with a composite summit caldera (Cioni et al., 1999), inside which the younger cone of “Gran Cono del Vesuvio” grew (Fig. 1). ⁎ Corresponding author. E-mail address: [email protected] (R. Scandone). 0377-0273/$ - see front matter © 2007 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2007.09.014

The volcano has erupted in wide range of styles. It is possible to distinguish different activity stages (Cioni et al., 1999): 1) 1st stage: Somma Volcano: predominant by effusive eruptions caused the early growth of the Somma, with only one reported major Plinian eruption (Codola eruption dated at 23–33 kA). 2) 2nd stage: Earlier caldera formation with at least 2 major Plinian eruptions (Basal Pumices, 17 ka, and Mercato, 8 ka) that caused two episodes of caldera formation and minor inter-plinian activity, possibly outside the caldera boundary. 3) 3rd stage: Transitional stage: Two major Plinian eruptions (Avellino Pumice, 3.5 kA, and Pompeii, 79 AD, enlarged the previous caldera but an increasing number of inter-plinian eruptions also occurred, mainly explosive between 3500 BP and 79 AD and then increasingly effusive since 79 AD, with 2 sub-plinian events (512 and 1631 AD) and a predominantly

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Fig. 1. Structural sketch map of Vesuvius (redrawn after Bianco et al., 1998).

effusive stage since 1631 (Cioni et al., 1999; Principe et al., 2004). During the 3rd stage, a long period of quiescence, starting possibly in 1139 AD (although uncertain eruptions are reported in 1500 and 1571, Guidoboni and Boschi, 2006), ended on 16 of December 1631 with a sub-plinian eruption. The eruption started a period of persistent activity, punctuated by violent strombolian eruptions that lasted for more than three centuries. Since April 1944, the volcano has entered a new quiescent stage (Carta et al., 1981). We try to understand the recent activity of Vesuvius within the context of mechanisms of ascent of magma to the surface. Ryan (1987) suggests that the ascent of magma is strongly controlled by a buoyancy mechanism. Scandone et al. (2007) suggest that magma ascends by injection of discrete batches of magma into a pre-formed crack network. After injection, the ascent velocity of a magma-filled crack is controlled by: 1) the pseudo-buoyancy force arising from the difference between the pressure gradients due to the density of the liquid and that of the surrounding rocks (Weertman, 1971; Pollard and Muller, 1976; Menand and Tait, 2002); 2) the stress field along the pathway (Nakamura et al., 1977; Shaw, 1980); 3) the physico-chemical properties of the magma. The abundance of volatiles may enhance the ascent by lowering the bulk density; on the contrary the loss of volatiles

during ascent may reduce the buoyancy, increasing both density and viscosity of ascending magma. Within this frame, both the density structure of the volcano and the confining stress place important constraints on magma ascent. The stratigraphy below Vesuvius is characterized by a sharp density transition at 2 km bsl from limestone (with a density of 2500–2700 kg/m3) to loose volcanic, alluvial and marine sediments (with an average density of 2300 kg/m3) (Bruno et al., 1998; Berrino et al., 1998). The sedimentary fill of the plain represents a low-density barrier to the ascent of dense magma and there is strong evidence that shallow magma accumulation below this barrier provides the source for extensive magma reservoirs, which fed the plinian and sub-plinian eruptions (Delibrias et al., 1979; Barberi and Leoni, 1980). The detailed reports of a relatively long eruptive period represent an invaluable dataset that can be used to understand not only the mechanisms of single eruptions, but also their temporal pattern. As many volcanic phenomena observed at Vesuvius in the past are presently observed at other volcanoes, its history may help to unravel the complexities of the life of other active volcanoes. In this paper we discuss the mechanism of persistent activity and violent strombolian eruptions based on a critical review of original historical sources of the period 1631–1944. We try to apply at Vesuvius the idea that small volume eruptions are controlled by the episodic arrival of single magma batches ascending because of a pseudo-buoyancy mechanism recently proposed for silicic volcanoes (Scandone et al., 2007). We then

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analyze the overall pattern of activity during the period 1631– 1944 to highlight the main changes in eruptive style that eventually caused the current prolonged quiescence that has persisted since 1944.

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effusive–explosive paroxysms is not always followed by a caldera collapse and repose as already observed by Arrighi et al. (2001). We propose a classification scheme slightly modified from that of Mercalli (1905) in which we distinguish:

2. The activity between 1631–1944 Mercalli (1883), Baratta (1897), and Alfano and Freidlaender (1929) proposed that the eruptive activity of Vesuvius between 1631 and 1906 could be explained by recurring cycles, with major eruptions representing the conclusion of a prolonged period of activity. More recently Arnò et al. (1987), and Carta et al. (1981) have used this concept of cycles to classify and study the activity in the period 1631–1944. The typical cycle statistically tested by Carta et al. (1981) is characterized by the sequence of: 1) A period of repose (generally not exceeding a few years) (R = repose); 2) A phase of mild strombolian activity with the building of a conelet within the crater (A = persistent activity), and, eventually, the emission of some lava flows (IE = intermediate eruptions) either within the crater or outside it; (Fig. 2) 3) A violent eruption (FE = final eruption) usually with a fast lava flow and strong explosions followed in several case by the formation of a small caldera and a new repose. However, Scandone et al. (1993) cast some doubts about the existence of cycles of activity. Arrighi et al. (2001) have shown that some explosive eruptions that were not followed by a repose interval had a violence that was comparable to that of eruptions considered to represent the end of a cycle. At the same time, eruptions with predominantly effusive activity may have volumes as large as those of mainly explosive eruptions. We conclude, therefore, that major eruptions do not represent the culmination of a period of activity, but that instead they represent casual occurrence of a particular set of boundary conditions that increases the probability of a more energetic eruption. We also observe that the occurrence of major

1) Normal strombolian activity (A); 2) Slow lava effusions (Slow Effusion = SE); 3) Violent strombolian eruptions (or paroxysm) (PE), are preceded by fast lava emission, and occasionally followed by periods of quiescence. Violent strombolian eruptions are characterized by lapilli and ash fall deposits resulting from eruptive columns some kilometers high (Walker, 1973, Arrighi et al., 2001, Valentine et al., 2006). We use also the more generic term of Paroxism in consideration of the mixed style of activity (fast emitted lava flow, followed by high lava fountains up to 1–2 km height). The main difference between the different states of activity is the eruption rate, which increases (non-linearly) from 1 to 3. Periods of repose may follow major paroxysms, but many shorter repose intervals may have gone undetected in the historical record. States 1 and state 2 may occur simultaneously and represent the predominant style of activity of the volcano that can persist for years. States 3 are episodic, lasting from hours to days, and are characterized either by major explosions lasting tens of minutes, or by rapid outflow of lava (reaching the lower part of the volcano in a few hours-days) eventually followed by sustained, high lava fountains and explosions. Depending on the eruption rate, such episodes may cause the formation of a collapse crater and be followed by periods of quiescence lasting up to 7 y. In other cases (e.g., the 1903 and 1929 eruptions) the paroxysms are not followed by any repose. No evident relationship exists between repose and erupted volume. Within this picture, it is possible to describe an early period, following the 1631 eruption, characterized mostly by the filling of the caldera by persistent activity and slow effusive eruptions. There are only scarce records of this period, except for the major explosive eruptions. Paroxysms are recorded in

Fig. 2. Gouache of Pietro Fabris of the Gran Cono with the small conelet as appeared in 1774 (Hamilton, 1774).

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1649, 1654, 1660, 1680, 1682, 1685, 1689 (Scandone et al., 1993). The first fast lava flow outside the crater along the flank of the volcano, followed by an explosive paroxysm, is reported in 1694 (Sorrentino, 1734). In total, Scandone et al. (1993) have accounted for 99 eruptions (their FE and IE) since 1631; 5 had a VEI of 3+(1737, 1779, 1794, 1822, 1906), and 12 had a VEI of 3. 53 such eruptions were accompanied by explosive paroxysms and are henceforth classified as PE (see Table A1 in the appendix for a summary of all the eruptions. A more detailed description of each eruption is found in Scandone et al. (1993). When viewed from a temporal perspective, it appears that PEs were numerous until 1872 (49 events); since that date, eruptions with only effusive activity (SE) have longer durations (Carta et al., 1981), and there has been a slow accumulation of lava, either on the flanks of the cone (building of several lava domes between 1872 and 1899) or, since 1913, filling the crater (and outpouring of lava from it). 3. The strombolian activity and slow effusions Persistent strombolian activity and slow effusions were the dominant styles of activity of Vesuvius between 1631 and 1944. The activity was characterized by mild explosions within the crater of Gran Cono, with the building of one or multiple scoria conelets (Fig. 1). This activity often occurred in association with slow lava effusion, which, depending on the topographic control of the crater, either remained confined in it, or were poured outside. Such activity was documented in numerous chronicles of scholars visiting the summit of the volcano and was the object of systematic observation since the building of Vesuvius Observatory in 1841. The slow emission of lava usually formed complex lava fields with pahohoe surfaces (e.g., 1820, 1858), or dome-like accumulations of many small flows in the proximity of the vent (1881–83, 1891–94, 1895–99). The last period of such activity between 1906 and 1944 was well documented because of the easy accessibility after construction of the Observatory road in 1872, and of a funicular railway in 1899. The complete detailed description of the period 1906– 1944 is reported only in Italian by Imbo' (1949). Here we summarize part of those observations and use them to obtain an

order of magnitude estimate of the average effusion rate for activity during this time period. The violent strombolian eruption of 1906 formed a crater within the Gran Cono, due to a combination of collapse within the feeding conduit and explosion cratering (Fig. 3). A crater volume was estimated at 84 × 106 m3 immediately after the eruption (Malladra, 1923). The crater was partly filled by landslides of the steep slopes, and in May 1913 its volume was reduced to 58 × 106 m3. On 10 May 1913, sudden subsidence of the SW bottom of the crater formed a funnel-shaped cavity 75 m deep and 150 m wide (Malladra, 1923). After 5 July 1913, lava outpouring partially filled the cavity and accompanied the formation of a small scoria cone constructed by mild strombolian activity. The cavity was entirely filled in July 1915 (Fig. 4). Assuming a total volume of lava equal to that initially occupied by the cavity (4.4 × 105 m3), we estimate a minimum effusion rate of 0.007 m3/s. Between 1915 and 1920 the bottom of the crater is raised by approximately 100 m with a volume of 12 × 106 m3 with an average effusion rate of 0.07–0.08 m3/s: an order of magnitude higher than the rate estimated for 1913– 1915. The first outpouring of lava outside the crater occurred on 28 November 1926. We hypothesize uniform filling of the crater between August 1920 and November 1926, and a minimum erupted volume of lava of 26 × 106 m3. The average effusion rate for this period is 0.13 m3/s, slightly higher than that of the previous period (Fig. 5). Between 1926 and 1944 the entire crater filled with lava (Fig. 4), thus adding 20 × 106 m3 to the erupted volume. However, outflows of lava also cover an area outside the crater of 3 × 106 m2 with an average thickness of 20–25 m (Imbo', 1949). Therefore, we estimate a total volume of 95 × 106 m3 during the period 1926–1944, with an average effusion rate of 0.17 m3/s. These estimates are lower by a factor of 3 than those of Imbo' (1949), which were based on daily eruption rates. We consider his value to be too high because his estimate would have resulted in a much earlier filling of the crater than was observed and more voluminous flows outside the crater. We may compare this effusion rate with other periods of slow effusion within or outside the crater, such as those that occurred

Fig. 3. The crater formed during the eruption of 1906 seen from the Somma (photo Perret, 1908 in Malladra 1923).

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Fig. 4. The cavity formed in May 1913 within the crater of the 1906 eruption was almost filled by lava in 1915 (photo Malladra, 1923).

between 1638–1694, 1858–1861, 1891–1894, 1895–1899. Below we consider each effusive period in sequence. The diameter of the caldera formed after the eruption of 1631 was estimated by contemporary scholars between half and one mile (Braccini, 1632), and an unknown depth. The volume must have been comparable with the volume of the erupted products, which ranges between 0.25 (including lithics) and 0.5 km3 DRE (Rosi et al., 1993; Rolandi et al., 1993). Persistent activity after the eruption started as early as 1638 (Alfano and Freidlaender, 1929). We hypothesize that the caldera was entirely filled as

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of 1694, when lava was first emplaced outside the crater (Sorrentino, 1734). We obtain an average effusion rate for this time period between 0.14 and 0.28 m3/s, the lower estimate being similar to that of the period 1913–1944. The second effusive period that we examine began on 27 May 1858, when a low effusion rate eruption initiated from vents at a height of 750 m asl on the SW flank of Vesuvius (Palmieri, 1895). The eruption lasted until April 1861. The complex lava field formed by this eruption did not extend very far from the vent, mostly filling a pre-existing deep creek (Vallone Grande) and the surrounding areas (Palmieri, 1895). The eruption started a period of unusually long SEs, each one lasting for years (Carta et al., 1981). The total volume erupted was of 120 × 106 m3 with an average effusion rate of 1.32 m3/s, about an order of magnitude higher than the crater filling phases. We suggest that the high effusion rate of the 1858 eruption is the result of the low elevation of the vent. This period was immediately followed by an eccentric eruption in 1861 that opened vents at a low altitude (250 m asl), with an unusually low volume of magma. The SEs of 1891–95 and 1895–99 were fed both by fractures propagating from the central crater and vents located at an average height of 850 m asl. During this period of activity, the central crater did not display any sign of activity, so we may regard these eruptions as representative of the crater filling phases. The volume erupted during the two episodes was 36 × 10 6 m 3 and 100 × 10 6 m 3 respectively (Alfano and Freidlaender, 1929), with average effusion rates of 0.38 and 0.75 m3/s. 3.1. Interpretation Strombolian activity, with persistent explosions and ejection of incandescent lapilli, requires that magma resides and degasses at a shallow depth within the volcanic edifice (as at Stromboli, Giberti et al., 1992). If we hypothesize a very long conduit (10–12 km), the low supply rate would make it difficult to maintain an open conduit for extended periods of time. As at Stromboli (Francalanci et al., 2004), it is plausible to

Fig. 5. The conelet in mild strombolian activity in the almost filled crater of Gran Cono as appeared in 1930 (from Imbo', 1949).

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hypothesize a shallow region where magma accumulates and feeds the strombolian activity or the slow effusion. In order to maintain continuous effusion, the shallow reservoir must be episodically refilled at a rate comparable to the amount of magma emitted by the strombolian activity or slow effusion eruption. The existence of a shallow accumulation zone is further supported by the lava composition. Most lavas erupted during the 1631–1944 period are strongly porphyritic, with 40–60% phenocrysts. The principal phenocrysts are clinopyroxene, leucite, and calcic plagioclase. The groundmass is holocrystalline and made of the same phases as the phenocrysts (Trigila and Benedetti, 1993). Compositional variations and trace element compositions point to a very homogeneous magmatic system over the entire period with a constant magma source and similar differentiation processes (Villemant et al., 1993). The depth of the shallow reservoir has been estimated by Fulignati et al. (2004) and Marianelli et al. (2005) through studies of melt inclusions in leucite and salitic pyroxenes in the lavas and basal scoriae of several violent strombolian eruptions (1723, 1944, 1794, 1822, 1872, 1906, 1944). They suggest a high aspectratio reservoir extending within a pressure range of 10 to 60– 90 MPa. Hypothesizing a density of 2500 kg/m3, we estimate a depth range between 400 and 2500–3600 m below the top of the volcano, which corresponds to the zone above the limestone basement (Brocchini et al., 2001). We suggest that this zone represents the area of expansion for magma-filled cracks where the confining stress is lower than in the carbonatic rocks. We can evaluate the expansion of magma-filled cracks from the elastic properties of the corresponding rocks. The halfthickness of a two-dimensional dike, very long in one in-plane dimension relative to the other, subjected to an ambient compressive stress σy perpendicular to the dike plane and an internal magma pressure P(x) is (Rubin, 1995): wð0Þ ¼

P  r0y G 1v

l¼

DP l M

ð1Þ

with: G = the shear modulus, ν is the Poisson ratio, and M the stiffness of the material. The order of magnitude of the elastic constants can be estimated through the velocities of propagation of seismic waves (Vp and Vs): G ¼ qVs2 E ¼ qVp2

ð2Þ

with E = 2G(1 + v) is the Young's modulus. Assuming average Vp and Vs of 4.1 and 3.6 km/s for the limestone and 3.1 and 2.6 km/s for the volcano-sedimentary fill, respectively (De Natale et al., 1998), we obtain the values shown in Table 1. Using these values, we estimate an expansion of 2.4 times of the maximum half-width of a magma-filled crack at the transition between limestone and alluvium. We suggest that the combination of low density of the volcano-sedimentary cover and its lower stiffness allow the accumulation at shallow depth of magma ascending through a deep fracture system. It is

Table 1 Elastic parameters of rocks Elastic parameters

Limestone

Alluvium

G E ν M

35 GPa 45 GPa 0.35 53 GPa

16 GPa 22 GPa 0.29 22 GPa

likely that the shallow reservoir has the shape of a thick dyke several meters wide that extends for about 2–3 km below the crater. 4. The violent strombolian eruptions (Paroxism) Violent strombolian eruptions at Vesuvius are characterized by a first phase with high eruption-rate lava flows advancing at high velocity, and reaching, in a few hours, the base of the volcano. A second phase usually follows, characterized by an even higher eruption rate and the formation of kilometers-high lava fountains or a sustained eruption plume up to heights of 10 km. The more energetic eruptions were followed by collapse of the Gran Cono, which may be further enlarged by the explosive activity. Some paroxysms caused small calderaforming events within the Gran Cono and occasional short quiescence periods. The collapses were sometime accompanied by phreatomagmatic explosions with the emission of fine ashes with abundant accretionary lapilli like in 1906 (Perret, 1924). The peculiar aspect of these eruptions is their sudden occurrence. In most cases, the days to weeks before the eruption, are characterized by only low levels of activity (Mercalli, 1905), except in a few cases (e.g., 1906). In several cases earthquakes or disturbances of the water table were observed a few days before the eruption; in many cases they heralded the opening of a lateral vent. The more common occurrence was an early surge of lava, which rapidly propagated outside the lower part of the Somma caldera, menacing the many villages at the base of the volcano (Torre del Greco was destroyed three times by such events). Differently from other volcanoes, the eruption rate at Vesuvius did not decline with time, but more often showed several pulses (e.g., 1872, 1906, 1944) that generated new lava flows advancing on the earlier cooling flows. Morphologically the lava field commonly had single channels and aa surfaces. The ensuing pulses may have had increasing eruption rates, which ultimately produced high lava fountains sustained for tens of minutes. At times, these fountains had an inclined trajectory that threw large ejecta kilometers from the vent. One such event in 1906 caused the collapse of a church roof in the village of S.Giuseppe and the death of more than 200 people. The explosive phases (Fig. 6) always occurred from the summit of “Gran Cono”, even when the early lava issued from flank vents (e.g., 1794, 1760, 1861). In many cases, but not always (e.g., 1804, 1805, 1929), the explosive phase was followed by collapse within the Gran Cono and phreatomagmatic explosions. During the 1944 eruption, the seismometer at Vesuvius Observatory recorded a seismic crisis coincident with this last phase.

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with the lava fountain is comprised between 90 and 150 m3/s. Similar estimates can be made for the eruptions of 1872 and 1906, which had a similar fountain heights. 4.1. Interpretation

Fig. 6. The gouache of Pietro Fabris of the lava fountain of 8 August 1779 drawn basing on a sketch taken during the eruption (Hamilton, 1779).

The volumes of magma erupted as lava flows during these eruption range generally between 5–20 × 106 m3 (Scandone et al., 1993, Nazzaro, 1997). During the eruption of 1872, the lava flow, which destroyed part of the villages of Massa and S. Sebastiano, had a volume of 20 × 106 m3 and was erupted in less than 2.5 days (Palmieri, 1895) with an average effusion rate of 92 m3/s. The eruption rate might have been even higher as there were two pulses with an interval of a few hours. During 1906, at the beginning of the eruption, there was a downward propagation of lateral vents toward the base of Gran Cono. Lava effusion lasted for 3.5 days, producing a volume of 20 × 106 m3 (Mercalli, 1906) at an average effusion rate of 66 m3/s. The lava flow destroyed the outskirts of Boscotrecase, and propagated so rapidly that it killed seven people who were not able to escape from their houses. Even in this case, there were two major pulses. The second one produced high lava fountains at the summit crater and caused a new outpouring of lava from vents higher on the flank of Gran Cono. Lava emission during the 1944 eruption lasted 3 days and 45 min with a total volume of 20 × 106 m3 (Imbo', 1949) and an average effusion rate of 76 m3/s. Imbo' reports that in the night of 20 March, 1944, the lava flow issuing on the northern flank had an eruption rate of 55 m3/s and velocity of 50–100 m/h. The estimate of the eruption rate for the lava fountains activity is more difficult. Scandone et al. (1986) made an estimate of 150 m3/s for the phase that marked the transition from lava fountains to a sustained eruption column during the 1944 eruption, so we may infer that the eruption rate associated

The scoriae erupted during the lava fountain phases of different violent strombolian eruptions range in composition from K-tephrites to K-phonotephrites; the groundmass composition is the result of mixing between a melt resident in a shallow reservoir and a more mafic melt (Marianelli et al., 2005). The dissolved volatile content in the melt inclusions of the mafic melt exceeds 5 wt.%, with H2O contents of melt inclusions in the range of 2.3 to 4.9 wt.% (Marianelli et al., 2005). Fulignati et al. (2004) find evidence that, at least during the 1944 eruption, the scoriae of the main layer of the lava fountain deposits resulted from the mingling between the shallow volatile-poor magma and the deep seated CO2–H2Orich magma. They suggest that the ascent of the gas-rich magma triggered the eruption. We agree with Fulignati et al. (2004) that episodic paroxysms are related to the arrival of volatile-rich magma batches; we further suggest that the entire dynamics of the eruption is controlled by the arrival of multiple, rapidly ascending batches. The fast ascent may prevent effective degassing of magma, thus permitting the retention of a higher volatile content and favoring higher buoyancy (Scandone et al., 2007; Gonnermann and Manga, 2007). The reason for the episodic, fast ascent of some batches is still unknown, but may be related to variation of stress boundary conditions as suggested by the recurrence rate of energetic PEs at Vesuvius, similar to that of tectonic earthquakes in the nearby Apennines (Marzocchi et al., 1993). We infer that the overpressure, exerted by the buoyant magma at the base of the shallow residing magma column, is transmitted throughout the fluid because of Pascal's Law. The immediate result is an increase in the magma flux, or in case of a choked flow, fracturing of the conduit walls by an action similar to that of an hydraulic hammer. At Vesuvius, as at Stromboli, this phase was commonly signaled by the sudden occurrence of harmonic tremor preceding the opening of lateral vents (e.g., the 1906 eruption of Vesuvius, or the 2007 eruption of Stromboli). The opening of lateral vents may cause a sudden drainage of the central eruptive column with an even higher magma discharge. After the start of the eruption, the arrival of distinct batches is testified by the impulsive character of effusive activity, with rapid drops of magma discharge followed by sudden renewals of the lava emission. If the impulses are sufficient to evacuate the shallow reservoir, the new batches may easily reach the surface with their full complement of volatiles and produce sustained lava fountains. Distinct pulses characterized the fountain phases (e.g., 1929, 1944, Signore, 1934, Imbò, 1949). If the shallow reservoir is not entirely filled with magma, isolated major explosions may occur without concomitant fast lava effusion (e.g., 1903). This recurrent eruption mechanism also characterizes activity at Stromboli, as has been recently observed during

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the eruption of February–April, 2007. On 15 March 2007, two weeks after the beginning of the eruption and after several impulsive lava outflows from a lateral vent, another impulse, observed and measured by a fixed infrared camera, preceded by 11 min the formation of a sustained lava fountain at the top of the volcano (Barberi et al., 2007). The products erupted in this phase were a highly vesicular and poorly crystalline pumice, representing a volatile-rich magma, mingled with a crystal-rich degassed magma similar to that of the lava flow. (Bertagnini et al., 2007). The rapid evacuation of magma residing in the shallow reservoir may result in the collapse of the upper part of the crater and formation of a small caldera. At the same time, the collapse of the hydrothermal system around the shallow reservoir into the hot zone occupied by magma can produce phreatic explosions that further enlarge the collapse structure. During this period of activity, all the phreatic explosions have been observed at the end of the eruptions in coincidence with collapse of the summit crater in 1779, 1794, 1822, 1906, and 1944 (Arrighi et al., 2001). We suggest that the phreatic explosions are likely caused by the ingression of water in the hot zone of drained magma. The main water table below the volcano is about 1600 m bsl, as estimated by the drilling of a geothermal exploratory well (Balducci et al., 1985), we may hypothesize that the paroxysms with final phreatic explosions drain magma to this depth, although we cannot exclude shallower water tables. The rather constant volume of the lava flows erupted during the violent strombolian events suggests that the shallow accumulation zone had a volume of 10–20 × 106 m3. During the major paroxysms this shallow reservoir must have been almost entirely evacuated, resulting in the subsequent deflation of the volcanic edifice (e.g., 1861). In Table 2, we report the volumes of magma erupted during some of the major paroxysms, the effusion rate of the lava flow, the repose (if any) after the eruption, and the inferred supply rate necessary to refill the evacuated shallow reservoir. This has been evaluated by the ratio between the erupted volume and the length of the following repose. Effusion rates range from 8–127 m3/s, with a median value of 29 m3/s. Estimated refilling rates are two to four orders of magnitude lower, ranging from 0.04–1.45 m3/s. Table 2 Parameters of selected violent strombolian eruptions Eruption (y)

Volume (×106 m3)

Lava effusion rate (m3/s)

Following repose (days)

Supply rate (m3/s)

1737 1760 1767 1779 1794 1822 1834 1850 1855 1868 1872 1906 1944

10 10 11 4 24 7 15 20 17 6 20 20 20

29 13 18 12 127 26 29 26 15 8 92 66 76

2704 1270 840 1495 557 592 119 1760 204 733 1325 2629 Continues

0.04 0.09 0.15 0.03 0.49 0.13 1.45 0.13 0.97 0.09 0.18 0.09 ?

The latter rates are similar to eruption rates inferred for SE activity. 5. Trends of activity The alternation between persistent activity, slow effusions and violent strombolian eruptions ended abruptly in 1944 with a prolonged quiescence of the volcano after the last paroxystic eruption. To understand this change of style, we analyzed statistically the entire period to determine whether significant changes of activity had occurred before the last quiescence. Carta et al. (1981) have already shown that the slow effusive eruptions became significantly longer after 1858; we tested whether similar changes occurred for the violent strombolian eruptions as defined in this paper, using the data from Scandone et al. (1993) catalogue. Although these data were originally divided into four categories, we performed our analysis on the violent strombolian eruptions characterized by a mixed explosive–effusive style. They occur episodically and we wanted to check if their rate of occurrence remained constant in time (e.g., a Poisson process). If a Poisson process controls the eruption occurrence, then the inter-event time intervals between successive eruptions should follow an exponential distribution. In order to verify this hypothesis, we used a standard method (see for example Marzocchi, 1996) that requires: 1) A qualitative test to check for trends in the number of eruptions; 2) If no trend is detected in the rate of occurrence, we search for autocorrelation of lengths between events; 3) If the series is not autocorrelated we test the possibility that the events are distributed according to a Poisson distribution. This method has been used to check the temporal distribution of the lateral eruptions of Etna (Salvi et al., 2006) and we refer to that paper for all the assumptions and technical aspects, which also apply here. We analyzed the temporal occurrence of violent strombolian eruptions by counting the number of events in a 30 y window since 1631 (Fig. 7). The catalogue is incomplete in the period 1631–1694 (Scandone et al., 1993), so we started our analysis in 1694. Fig. 6 show a decreasing trend in the number of eruptions with time until 1944; superimposed on the histogram is a linear fit to the data: ŷ =A +Bx. The values of A, B and R are shown in the legend. In order to verify the goodness of the fit we performed a t-test for the slope. The null hypothesis H0 is that the angular coefficient B = 0 (no trend in the data). We reject the null hypothesis at a significance level of 0.002. We also checked for the autocorrelation of the length of intervals and rejected the hypothesis that the data are not correlated (highly significant correlation). Since the eruption data show a decreasing trend versus time and are autocorrelated, we cannot assume that they are described with a homogeneous Poisson process. In order to get a statistical evaluation of eruption probabilities in the presence of a trend, we must look for a generalization of the

R. Scandone et al. / Journal of Volcanology and Geothermal Research 170 (2008) 167–180

175

Fig. 7. Histogram and linear fit of the number of violent strombolian eruptions in the period 1694–1944. The values on the vertical axis are the number of eruptions occurring in a time window of 30 y.

Poisson process. In the Poisson process, the probability of observing R(= r) events is: Pð R ¼ rÞ ¼ ek

kr r!

ð3Þ

with λ = constant the average number of events per unit of time. Eq. (3) can be generalized by allowing λ to vary as an increasing or decreasing function of time. Substituting λ with the time-dependent function λ(t,θ), we obtain another kind of Poisson process named a non-homogeneous Poisson process (Ho, 1996). In this case, the intensity function λ(t,θ) (θ defined in the following) is the instantaneous rate of change of eruptions with respect to time and is called the instantaneous recurrence rate of the volcanic process. Here we model the volcanic time-trend using a power law process, letting t be predetermined and supposing that n N 1 eruptions are observed during [0,t] at time 0 b t1 b t2 b … b tn bt.The intensity function (instantaneous recurrence rate) is b
t b1 kðt; hÞ ¼ ð4Þ h h at time t (θ is parameter, and β is a coefficient of shape; when β is equal to 1, the time dependence of λ vanishes). Ho (1996) and Crow (1974) use the Maximum Likelihood Method to estimate θ, and β: defining: S¼

n X

ln ðt=ti Þ

ð5Þ

i¼1

n bˆ ¼ S hˆ ¼

ð6Þ t ˆ

n1= b where β̑ and θ̑ are the estimators of β and θ.

ð7Þ

In the application of the power law process to volcanic eruptive forecasting, the estimate of λ(t,θ) is of considerable practical interest because it represents the instantaneous eruptive status of activity at the end of the observation time t. By using the estimators (5) and (6), the estimate of (4) is:  ˆ bˆ t b1 n bˆ : ¼ kˆðt; hÞ ¼ t hˆ hˆ

ð8Þ

The power law process generalizes the homogeneous Poisson process, because when β = 1, the power law process reduces to a homogeneous Poisson process. When β N 1, the sequence has an increasing number of events, and when β b 1 has a decreasing number of events. We obtain the following values: b ¼ 0:7494; h ¼ 1:5628; kðt; hÞ ¼ 0:1343 for n = 45, and t = 250 y. The violent strombolian eruption data set has β b 1, due to the decreasing number of paroxistic eruptions with time. These findings suggest that there is a decreasing probability of occurrence of a violent strombolian eruption. We estimate that the yearly probabilities of having at least one eruption in 1944 and in 2007, given by 1  Pð R ¼ 0Þ ¼ 1  ek are respectively 0.12 and 0.10, with λ estimated for 250 and 313 y respectively. We cannot make a similar estimate of the probability of having an eruption of different magnitude (e.g., sub-plinian or plinian) for the limited Vesuvian database. 6. Discussions and conclusions We suggest that the entire 1631–1944 period is characterized by non-stationary strombolian activity with magma residing

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high in the volcanic edifice, but that there was a progressive change of the style of activity with longer times spent in slow effusive activity and progressively fewer violent strombolian eruptions with time. The residence of magma at shallow depth is favored by a relatively denser, shallow inner core of the volcano made by a network of cooled dykes (Zollo et al., 1996). The denser core causes a high level of the neutral buoyancy (Ryan, 1987). Periodic input of magma into the shallow accumulation zone causes fluctuations of the height of the magma column and occasional overflow, with slow effusion of magma either within the crater or outside of it. The eruption rate of slow overflows is representative of the “normal” magma supply of the volcano. Slightly higher eruption rate associated with SEs from lower vents can be explained with the higher buoyancy forces that drive these eruptions. The magma supply rate, as evaluated by slow eruptions (SE) or caldera filling, varies by up to two orders of magnitude, from ∼0.01–1 m3/s. We hypothesize that the repose following some violent strombolian eruptions is the time required to fill the shallow accumulation zone, thus the ratio of erupted lava/repose time may be used to estimate the magma supply rate. Table 3 reports a summary of all the supply rates evaluated by different methods. As higher supply rates result from shorter times, we evaluate a time-weighted average supply rate of 0.18 m3/s, which covers about 140 of the 313 y between 1631 and 1944. The wide variations from the mean suggest that this value represents the time-averaged arrivals of slow magma batches into the shallow reservoir. In contrast to the supply rate, the lava effusion rates of the paroxistic eruptions have less variations and similar eruption dynamics, even for events occurring from eccentric vents like those of 1760, 1794, and 1861. Table 3 Estimates of supply rates Time

Duration (days)

Supply rate (m3/s)

Method

1638–94 1737 1761 1767 1779 1794 1822 1834 1850 1855 1858–61 1868 1872 1891–94 1895–99 1906 1913–15 1915–20 1920–26 1926–44

22069 2704 1270 840 1495 557 592 119 1760 204 1051 733 1325 1092 1527 2629 730 1825 2190 6570

0.13–0.26 0.043 0.089 0152 0.025 0.488 0.129 1.45 0131 0.965 1.32 0.094 0.175 0.38 0.75 0.088 0.007 0.07 0.13 0.17

Crater fill Repose Repose Repose Repose Repose Repose Repose Repose Repose Slow effusion Repose Repose Slow effusion Slow effusion Repose Crater fill Crater fill Crater fill Crater fill

The highly porphyritic nature of erupted lavas indicates that the average supply rate does not balance the heat loss to the surrounding rocks so that magma cannot reside without appreciable cooling, or else the persistent activity and slow effusion cannot maintain a stationary state of activity. Alternatively, crystallization caused by degassing requires that the shallow reservoir be above the H2O saturation level and loss of volatiles causes enhanced crystallization. Periodic impulses of rapidly ascending magma served to rid the upper conduit of the cooler and crystal-rich magma, representing a sort of cleaning of the upper part of the volcano. This mechanism was in operation for 313 y between 1631 and 1944. However, the progressive decrease in the number of violent strombolian eruptions with time gradually made this process less effective. The volcano finally entered a quiescent stage because the shallow plug crystallized extensively before the arrival of the next energetic magma batch. We do not know if the decrease in the number of violent strombolian eruptions has been caused by the accumulation of a dense crystal mush in the shallow reservoir, which acted as a more effective barrier to the ascent of fast batches, or if it reflects an actual change of the state of stress of the volcano. There must be a substantial accumulation of magma along the ascent path of the volcano since 1944, if the deep supply of magma has not changed. We reach, by different arguments, the same conclusions as Santacroce (1983) regarding the storage of magma at depth, although we believe that it is difficult to estimate the amount of eruptible magma because we do not know the actual magma supply and the cooling and crystallization rate. We suggest that during the period 1631–1944 1) The deeper part of the magma supply system always fed the shallow accumulation zone, thus suggesting a preferential ascent path along the same fracture zone; 2) The ascending magma batches acted as a piston that pushed up the resident magma with limited mingling; this suggests a limited width of the shallow reservoir, possibly that of a dyke 2–3 km deep. 3) The ascent of fast magma batches during paroxystic eruptions was not signaled by seismicity, except when the magma flux could not be accommodated by the existing shallow conduit and caused the opening of lateral vents or spreading of the subsurface structure of the volcano. 4) It is likely that the NE–SW regional fault crossing Vesuvius, evidenced at sea by seismic profiles (Finetti and Morelli, 1974), represents the main direction of the dyke feeding the volcano; along this lineament are located several eccentric vents, including the lateral one that fed the reawakening of the volcano in 1631. 6.1. The present The 1944 eruption emitted a volume of lava similar to that of previous events, and ended with phreatomagmatic explosions and the collapse of Gran Cono (Fig. 8). It is likely that the

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177

Fig. 8. The “Gran Cono del Vesuvio” after the eruption of 1944; in the background the northern lava flow of the 1944 eruption and the Somma caldera wall (photo Scandone).

shallow system was substantially evacuated. The first significant anomalies were observed from 1952 to 1954 with an increase in the temperature of fumaroles within the crater from 350 to 600°C (Imbo' et al., 1964a), and on 11 May, 1964, with an earthquake swarm with a widely felt shock (Imax = 5 Mercalli Scale) (Imbo' et al., 1964b). Since then, the temperature of the fumaroles has decreased (Nazzaro, 1997). No major seismic crises, such as those that might be associated with an impulsive arrival of rapidly ascending magma, occurred until 1978–80, when there was a swarm of moderate earthquakes (M b 3). The most significant earthquake since 1964 occurred in 1999 with Md = 3.6, accompanied by a relatively small number of earthquakes with Md N 1.8. Since 1985–86, there has been a progressive decrease of the b-value of the earthquakes (the measure of the relationship between the number and magnitude of earthquakes in the Gutenberg– Richter law), an alternating sequence of seismic quiescence and activity, and an increase in the observed earthquake magnitude (De Natale et al., 2004). The source mechanisms of the major earthquakes are compatible with isotropic components that indicate volumetric expansion (De Natale et al., 2004). The recurrence rate of seismic swarms or thermal anomalies in the past 50 y is =0.14 y− 1, similar to the one for paroxistic eruptions evaluated until 1944 (0.13 y− 1). Collectively these data suggest that the upper feeding system is being intruded by new episodic arrivals of magma batches that are progressively entering in a shallow accumulation zone. The surface manifestation of this activity is an overall spreading of the volcano with a subsidence of the Somma Caldera. SAR images give evidence of a regional trend, characterized by eastward movement of the SE sector relative to the NW sector, with the transition between W and E displacements along a NE–SW alignment passing through the northern flank of Vesuvius (Borgia et al., 2005).

Scandone et al. (2007) suggested that the episodic ascent of discrete magma batches explains the dynamics of small eruptions at silicic volcanoes. We have shown that the same model can equally well explain the violent strombolian eruptions of Vesuvius and possibly of other basaltic volcanoes with persistent activity. Valentine and Krogh (2006) evidenced in the Paiute Ridge (Nevada, USA) that dykes mostly occupy pre-existing normal faults eventually forming shallow sills at a few hundreds of meters from the paleo-surface. We believe that the overall trend of activity of volcanoes is strongly controlled by the interplay between the ascent mechanism of magma and its storage in crustal reservoirs, and that our predictive capability depends mostly on the knowledge of the previous volcanic history of the volcano as well as its inner structure. We surmise that a possible re-awakening of the volcano might be triggered by new arrivals of rapidly ascending magma batches with enough energy to intrude the cooler shallow system; if so, it is possible that renewal of activity may occur along the same zone of weakness as that of 1631, namely along the NE–SW fault zone. It is likely that new episodic seismic crises with shallow hypocenters, even of short duration, may accompany these intrusions and the eventual outbreak at the surface. Acknowledgements We deeply acknowledge Kathy Cashman for the critical reading of an earlier version of the manuscript. Greg Valentine and Patrick Bachélerie made useful suggestions for the improvement of the manuscript. We acknowledge partial financial support from INGV and Dipartimento Protezione Civile Nazionale of Italy, and MIUR-Prin Project 2005 (Risalita dei Magmi e Dinamica delle Eruzioni).

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Appendix A

Table (continued) A1 (continued )

Table A1 List of eruptions of Vesuvius between 1631 and 1944

Beginning (d-m-y)

End (d-m-y)

12-8-1805 31-5-1806 4-9-1809 11-9-1810 1-1-1812 9-10-1813 29-10-1813 22-12-1817 1-12-1819 15-1-1822 21-10-1822

19-10-1805 9-6-1806 5-9-1809 22-9-1810 28-2-1812 28-10-1813 28-2-1814 26-12-1817 31-5-1820 28-2-1822 16-11-1822

68 9 1 11 58 19 122 4 182 44 26

Beginning (d-m-y)

End (d-m-y)

Length (days)

15-12-1631

2-1-1632

18

28-11-1649 25-2-1654 3-7-1660 26-3-1680 12-8-1682 3-10-1685 9-12-1689 13-4-1694 31-7-1696 15-9-1697 19-5-1698 1-7-1701 19-5-1704 29-7-1707 5-2-1712 25-10-1712 12-4-1713 6-1-1714 15-6-1714 6-6-1717 22-12-1717 3-9-1718 7-5-1720 1-5-1721 20-4-1723 4-9-1724 22-4-1727 27-2-1730 25-12-1732 1-7-1735 14-5-1737

10-3-1650 1-3-1654 29-7-1660 28-3-1680 22-8-1682 10-10-1685 16-12-1689 29-4-1694 14-8-1696 9-1-1698 15-7-1698 15-7-1701 23-5-1704 22-8-1707 10-6-1712 8-11-1712 25-5-1713 20-1-1714 30-6-1714 18-6-1717 26-12-1717 9-7-1719 27-5-1720 7-6-1721 8-7-1723 29-9-1724 27-4-1727 1-4-1730 10-1-1734 30-7-1735 4-6-1737

102 4 26 2 10 7 7 16 14 116 57 14 4 24 126 14 43 14 15 12 4 309 20 37 79 25 5 33 381 29 21

25-10-1751 2-12-1754 29-3-1759 13-5-1759 6-11-1759 23-12-1760

25-2-1752 15-3-1755 31-3-1759 20-5-1759 30-3-1760 6-1-1761

123 103 2 7 145 14

28-3-1766 19-10-1767

15-12-1766 27-10-1767

262 8

15-2-1770 1-5-1771 29-12-1773 4-8-1774 20-12-1775

30-4-1770 30-5-1771 1-2-1774 1-12-1774 3-4-1776

74 29 34 119 105

29-7-1779

13-8-1779

15

1-7-1785 1-8-1788 5-9-1790 16-6-1794

30-11-1787 15-8-1788 16-11-1790 5-7-1794

882 14 72 19

12-8-1804

28-11-1804

108

Type Volume

VEI

PE Volume pyroclasts = 0.2 km3 PE PE PE PE PE PE PE PE SE SE PE PE PE PE SE SE PE PE PE PE SE SE PE SE PE PE PE PE SE SE PE Volume lava = 10 × 106 m3 PE PE PE SE SE PE flank Volume lava = 9.8 × 106 m3 PE PE Volume lava = 11 × 106 m3 pyroclasts 1.4 × 106 m3 PE PE SE SE SE Volume lava = 48.7 106 m3 PE Volume lava = 3.5 × 106 m3 pyroclasts = 6.1 × 106 m3 PE PE PE PE flank Volume lava = 23.5 × 106 m3 PE

4

2

Length (days)

3 2 3 2 2 3 2 2 3 2 1 3 1 1 1-2 1 2 2-3 1 2 1

14-8-1831 25-5-1833 27-11-1833 22-8-1834

23-12-1832 15-6-1833 16-1-1834 2-9-1834

497 21 50 11

1-1-1839 5-2-1850

3-1-1839 16-2-1850

2 11

1-5-1855

27-5-1855

26

27-5-1858

12-4-1861

1051

8-12-1861 15-11-1867 15-11-1868

31-12-1861 31-5-1868 26-11-1868

23 198 11

13-1-1871 24-4-1872

30-4-1871 30-4-1872

107 6

3 2 2 2-3 2 1 3–4

16-12-1881 2-5-1885 16-4-1887 1-5-1889 7-6-1891

31-1-1884 1-7-1886 19-4-1887 30-9-1889 3-6-1894

776 425 3 152 1092

3-7-1895

7-9-1899

1527

27-8-1903

30-9-1904

400

1 1 3

3-2-1906

3-4-1906

59

4-4-1906

22-4-1906

18

1 3

27-11-1926 1-8-1927 11-8-1928 3-6-1929

28-11-1926 2-8-1927 12-8-1928 8-6-1929

1 1 1 5

11-7-1930 2-10-1930 1-6-1933 12-2-1935 8-7-1935 28-3-1936 4-6-1937 8-8-1939 26-6-1940 22-10-1941 6-1-1944 18-3-1944

30-7-1930 9-11-1930 19-11-1934 31-3-1935 21-8-1935 24-9-1936 7-7-1937 9-8-1939 31-7-1940 15-12-1942 23-2-1944 4-4-1944

19 38 536 47 44 180 33 1 35 419 48 17

2 2

2 2 1 1 1–2 3–4

1 1 3–4 2

Type Volume

VEI

PE PE SE PE SE SE PE SE SE SE PE Volume lava = 7 × 106 m3 SE SE SE PE Volume lava = 15 × 106 m3 PE PE Volume lava = 20 × 106 m3 PE Volume lava = 17 × 106 m3 SE volume lava = 120 × 106 m3 PE Lateral PE PE Volume lava = 6 × 106 m3 SE PE Volume lava = 20 × 106 m3 SE SE SE SE SE Volume lava = 36 × 106 m3 SE Volume lava = 100 × 106 m3 PE Volume lava = 0.5 × 106 m3 SE Volume lava = 1.8 × 106 m3 PE Volume lava = 20 × 106 m3 SE SE SE PE Volume lava = 12 × 106 m3 SE SE SE SE SE SE SE SE SE SE SE PE Volume lava = 20 × 106 m3

2–3 2 1 1 1 1 2 1 1–2 2 3–4 1–2 1–2 1 3 2–3 2–3 2–3 2 3 1 2–3 1 3 1–2 1 1 1 2 2 2 1 3–4 1 1 1 2–3 1 1 1–2 1 1– 1 0 0 0 1–2 0 3

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