Journal of Volcanology and Geothermal Research 134 (2004) 77–107 www.elsevier.com/locate/jvolgeores

A discussion of the mechanisms of explosive basaltic eruptions

Elisabeth A. Parfitt

Department of Environmental Science, Lancaster University, Lancaster, LA1 4YQ, UK Received 3 February 2003; accepted 16 January 2004

Abstract

Two contrasting models of the dynamics of explosive basaltic eruptions are in current usage. These are referred to as the rise speed dependent (RSD) model and the collapsing foam (CF) model. The basic assumptions of each model are examined, and it is found that neither model is flawed in any fundamental way. The models are then compared as to how well they reproduce observed Strombolian, Hawaiian and transitional eruptive behaviour. It is shown that the models do not differ greatly in their treatment of Strombolian eruptions. The models of Hawaiian eruptions are, however, very different from each other. A detailed examination of the 1983–1986 Pu’u ‘O’o eruption finds that the CF model is inconsistent with observed activity in a number of important aspects. By contrast, the RSD model is consistent with the observed activity. The issues raised in the application of the CF model to this eruption draw into doubt its validity as a model of Hawaiian activity. Transitional eruptions have only been examined using the RSD model and it is shown that the RSD model is able to successfully reproduce this kind of activity too. The ultimate conclusion of this study is that fundamental problems exist in the application of the CF model to real eruptions. D 2004 Elsevier B.V. All rights reserved.

Keywords: basaltic; explosive; eruption; strombolian; hawaiian; foam; separated flow

1. Introduction almost certainly to eruptions on Mars (Wilson and Head, 1983, 1994). The presence of dissolved gas Basaltic volcanism is the dominant mode of vol- within basaltic magma results in explosive volcanic canic activity on Earth, the Moon, Mars and Venus activity unless the exsolution of the gas from the (e.g., Cattermole, 1989; Head et al., 1992; Wilson and magma is suppressed (as in sufficiently deep sea-floor Head, 1994). On Earth, >80% of the annual volcanic volcanism—Head and Wilson, 2003) or the gas is lost output is basaltic with >70% of this occurring beneath from the magma prior to eruption. Although explosive the Earth’s oceans (Crisp, 1984). Basaltic eruptions basaltic eruptions are generally much less violent than are frequently described as effusive because they their more silicic counterparts they are, nonetheless, commonly generate lava flows. While the term ‘‘ef- explosive and need to be considered as part of a fusive’’ is appropriate for basaltic eruptions in which continuum of explosive activity that embraces not the lava oozes passively from the vent, it is a mis- only the familiar explosive basaltic eruption styles— leading term when applied to the majority of subaerial Hawaiian and Strombolian—but includes sub-Plinian, eruptions on Earth, to eruptions on the Moon and Plinian, ultra-Plinian and ignimbrite-forming events. Our understanding of the mechanisms of explosive basaltic eruptions has advanced considerably during E-mail address: [email protected] (E.A. Parfitt). the past f30 years due to the collection and analysis

0377-0273/$ - see front matter D 2004 Elsevier B.V. All rights reserved. doi:10.1016/j.jvolgeores.2004.01.002 78 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 of new field data (e.g., Heiken, 1972, 1978; Walker, al., 1986; Bertagnini et al., 1990). Though rare exam- 1973; McGetchin et al., 1974; Self et al., 1974; Self, ples of sub-Plinian and Plinian basaltic activity do 1976; Williams, 1983; Walker et al., 1984; Houghton occur (Self, 1976; Williams, 1983; Walker et al., and Schmincke, 1989; Carracedo et al., 1992; Thor- 1984), explosive basaltic eruptions resulting from darson and Self, 1993; Parfitt, 1998), volcano moni- the exsolution of magmatic gases alone (rather than toring (e.g., Richter et al., 1970; Chouet et al., 1974; hydromagmatic activity) generally exhibit Hawaiian Blackburn et al., 1976; Swanson et al., 1979; Wolfe et or Strombolian styles, or behaviour which exhibits al., 1987, 1988; Neuberg et al., 1994; Vergniolle and characteristics of both end-member styles. Brandeis, 1994, 1996; Ripepe, 1996;Vergniolle et al., 1996; Hort and Seyfried, 1998; Chouet et al., 1999), 2.1. Hawaiian activity laboratory studies (e.g., Jaupart and Vergniolle, 1988; Mangan et al., 1993; Mangan and Cashman, 1996; The term ‘‘Hawaiian’’ is used to denote eruptions Zimanowski et al., 1997; Seyfried and Freundt, 2000) that are continuous in character and that generate lava and through mathematical modelling (Sparks, 1978; fountains (Fig. 1), typically tens to hundreds of metres Wilson, 1980; Wilson and Head, 1981; Stothers et al., in height (though they can occasionally exceed 1 km 1986; Vergniolle and Jaupart, 1986; Head and Wilson, in height: Wolff and Sumner, 2000). As the term 1987; Jaupart and Vergniolle, 1988; Woods, 1993; suggests, this type of activity is characteristic of the Parfitt and Wilson, 1995, 1999). It is now widely volcanoes of the Hawaiian chain but it is commonly accepted that Strombolian eruptions result from the seen on other basaltic volcanoes, e.g., Eldfell (Self et formation and bursting of a gas pocket close to the al., 1974), Hekla (Thorarinsson and Sigvaldason, surface (e.g., Blackburn et al., 1976; Wilson, 1980; 1972), Etna (Bertagnini et al., 1990) and Miyakejima Vergniolle and Brandeis, 1994, 1996), though some (Aramaki et al., 1986). Hawaiian eruptions have details of the mechanism are still disputed and are typical durations of hours to days, during which time discussed below. In the case of the dynamics of a lava fountain of fairly constant height may play Hawaiian eruptions, however, a curious situation above the vent (e.g., Wolfe et al., 1988). The lava exists in which two very different models have been fountain ejects clots of magma ranging in size from developed that are both in common usage. I refer to millimetres to about a metre in diameter into the air at 1 these models as the rise speed dependent (RSD) speeds of typically f100 m sÀ (Wilson and Head, model (Wilson, 1980; Wilson and Head, 1981; Head 1981). The majority of the erupted material lands and Wilson, 1987; Fagents and Wilson, 1993; Parfitt and Wilson, 1994, 1999; Parfitt et al., 1995) and the collapsing foam (CF) model (Vergniolle and Jaupart, 1986, 1990; Jaupart and Vergniolle, 1988, 1989; Vergniolle, 1996). The aims of this paper are to review both models of explosive basaltic eruptions, and to present an in- depth examination of the models of Hawaiian activity in which the assumptions and predictions of each model are compared with a wide range of geophysical and observational data from recent eruptions.

2. Styles of explosive basaltic eruption

Volcanologists have had many opportunities to observe and monitor explosive basaltic eruptions Fig. 1. Photograph of a lava fountain at the Pu’u ‘O’o vent. The (e.g., Richter et al., 1970; Blackburn et al., 1976; fountain is f400 m in height. (Photograph taken by Lionel Wilson, Swanson et al., 1979; Fedotov et al., 1983; Aramaki et August 1984). E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 79 while still incandescent, and accumulation and coa- produced a number of lava flows, the longest of which lescence of these hot clots generates rootless lava reached a length of 27 km (Lockwood et al., 1987). flows (Head and Wilson, 1989). These flows are Much material falling from the outer edges of the typically still fluid enough to flow many kilometres fountain cools sufficiently during flight that, though it to tens of kilometres from the vent. For example, a 21- deforms on landing and is hot enough to weld to the day-long Hawaiian eruption at Mauna Loa in 1984 material around it, is not hot enough to form rootless

Fig. 2. Hot clots of magma accumulate around vents forming spatter ramparts/cones. (a) A section of a spatter rampart formed during the April 1982 eruption of Kilauea. Individual clots have flattened and flowed upon landing. Each clast is f0.2 m is diameter and is welded to those above and below them. (Photograph taken by the author). (b) The spatter cone and down-wind tephra blanket formed during the 1959 Kilauea Iki eruption. Close-up the cone is formed of welded clasts like those in (a). The figure is standing in a collapse pit within the down-wind tephra blanket. Here, at a distance of f0.5 km from the vent, the deposit is composed of centimetre-scale clasts and is unwelded. (Photograph taken by the author, May 1996). 80 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 lava flows and instead accumulates as a spatter cone around the erupting vent (Fig. 2; Head and Wilson, 1989). Some even cooler material can accumulate to form a loose cinder cone, and a small proportion of the erupted material is carried upwards in a convective plume above the fountain and is deposited downwind forming a tephra blanket (Fig. 2b, Parfitt, 1998).

2.2. Strombolian activity

Strombolian activity takes its name from the fre- quent, small-scale, transient explosions exhibited by Stromboli, a volcano which forms one of the Aeolian Islands north of Sicily. Whereas the term ‘‘Hawaiian’’ is well-defined and used in a fairly restricted way, the term ‘‘Strombolian’’ has been used to denote a wide range of activity, and, thus, caution must be used in understanding individual usage of the term. The term ‘‘Strombolian’’ is most commonly used (and is used here) to denote the relatively mild explosions that occur from the accumulation of gas beneath the cooled upper surface of a magma column (e.g., Black- burn et al., 1976; Wilson, 1980). In such events, gas accumulation causes a raising and up-doming of the surface of the magma column. This ‘‘blister’’ eventu- Fig. 3. Photograph of the plume generated during an explosion at ally tears apart allowing the release of the gas and the Stromboli. The plume is f200 m in height. (Photograph taken by ejection of the magma that formed the skin of the the author, September 1996). blister. Blackburn et al. (1976) found typical initial 1 velocities of clasts at Heimaey to be f150 m sÀ whereas at Stromboli initial velocities are generally Though many Strombolian explosions are mild, 1 50–100 m sÀ (Chouet et al., 1974; Blackburn et al., discrete events, the term Strombolian is also used to 1976; Weill et al., 1992; Vergniolle and Brandeis, describe events which can generate sustained eruption 1996). Each explosion usually lasts f1 s and one plumes that reach heights of up to 10 km above the explosion may follow another after anything from a vent (e.g., Cas and Wright, 1988). These are events in few seconds to several hours. At Stromboli the typical which the individual explosions are so closely spaced time between explosions is between 10 min and 1 in time that they generate a sustained eruption plume h (Vergniolle and Brandeis, 1996). The erupted ma- of considerable height rather than the small plumes terial is generally cooler prior to eruption than that associated with truly discrete explosions (e.g., Fig. 3). produced in Hawaiian eruptions and also experiences The 1973 Heimaey eruption in Iceland provides a more cooling during flight than Hawaiian clasts. The good example of this type of behaviour. The eruption clasts produced are too cool on landing to weld or produced explosions every 0.5–3 s with incandescent coalescence and so accumulate as a tephra/cinder cone clasts reaching heights of f250 m above the vent and around the vent (McGetchin et al., 1974; Heiken, generated a plume that extended to heights of 6–10 1978). At Stromboli clasts typically reach heights of km above the vent (Self et al., 1974; Blackburn et al., <100 m above the vent (Vergniolle and Brandeis, 1976). The eruption simultaneously generated lava 1996) and the plume generated by the explosion flows. This behaviour is distinctly different from the generally rises to heights of <200 m (Fig. 3,J. discrete explosive events seen at Stromboli and Davenport, unpublished data). appears, in fact, to represent a type of behaviour E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 81 which exhibits characteristics of both Hawaiian and 3. Models of eruption mechanisms Strombolian eruptions: Although the explosions are discrete, they are so closely spaced in time that in 3.1. The rise speed dependent model terms of the eruption column the activity is continuous as in a Hawaiian eruption. The continuous production The earliest attempt to apply fundamental ideas of of lava flows is also more characteristic of Hawaiian conservation of energy and mass in volcanic eruptions events than the mild Strombolian events described was made by McGetchin and Ulrich (1973), but they previously. Thus, this type of eruption can be viewed applied their model only to eruptions producing maars as transitional between the Hawaiian and Strombolian and diatremes. The first model to specifically address end-member eruption styles. the dynamics of explosive basaltic eruptions was Classification schemes for explosive basaltic activ- developed by Wilson (1980) and Wilson and Head ity define Strombolian events as being more ‘‘explo- (1981). These two papers set out the basic premises of sive’’ than Hawaiian events (Fig. 4; Walker, 1973; Cas the RSD model that have been developed further in and Wright, 1988). Two points are important to note subsequent papers (Head and Wilson, 1987; Fagents about such classification schemes: (1) they are based and Wilson, 1993; Parfitt and Wilson, 1994, 1995, on the dispersal in eruptions like the Heimaey erup- 1999; Parfitt et al., 1995). The essential idea set out in tion, not on truly discrete Strombolian explosions like these papers is that Strombolian and Hawaiian activity those occurring with such regularity at Stromboli; and represent end-members of a continuum of explosive (2) they can lead to misclassification of Hawaiian basaltic activity and that the form of activity that eruption deposits. The 1959 Kilauea Iki deposits, for occurs depends most fundamentally on the rise speed example, would be classified as Strombolian (Fig. 4) in of the magma beneath the eruptive vent (e.g., Table 1). such a scheme when they were actually deposited Volatiles exsolve from magma as it rises, and the during a classic Hawaiian eruption (Richter et al., depth at which exsolution occurs depends on the 1970; Parfitt, 1998). Thus, it is important to exercise volatile species and the amount of dissolved volatiles caution in the use of the terms Hawaiian and Strombo- present (Wilson and Head, 1981). Gas bubbles that lian and to recognise that they represent end-member form within the magma are always buoyant and rise cases while many basaltic eruptions simultaneously upwards through the magma at a rate that depends on exhibit facets of both types of activity and are better the size of the bubble and the magma rheology. In the described as ‘‘transitional’’ eruptions (Parfitt and Wil- RSD model, it is assumed that if the rise speed of the son, 1995). magma is relatively great then the bubbles do not rise

Fig. 4. Diagram showing Walker’s (1973) classification scheme for explosive volcanic eruptions which is based on the degree of fragmentation (F) of the magma and the dispersal area (D) of the tephra. The asterisk shows that the deposits of the 1959 Kilauea Iki eruption would be classified as Strombolian using this scheme even though the deposits were formed during a typical Hawaiian eruption. Redrawn from Cas and Wright (1988). 82 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107

Table 1 Rise speeds beneath vents during recent Hawaiian (H), Strombolian (S) and transitional (T) eruptions Eruption Style Volume flux Vent area Rise speed at Reference 3 1 2 1 (m sÀ ) (m ) depth (m sÀ ) Mauna Ulu, 1969 S 0–3 400 0–0.008 Swanson et al. (1979) 3 3 Stromboli, 1971 S 8 10À 2.3 3.5 10À Chouet et al. (1974) Â Â Kupaianaha, 1986 S 3 315 0.01 Parfitt and Wilson (1994) Etna, typical Strombolian activity S <9.1 – 0.0006–0.045 Harris and Neri (2002) Heimaey, 1973 T* 30 314 0.1 Self et al. (1974), Blackburn et al. (1976) Kilauea Iki, 1959 H 160 180 0.9 Richter et al. (1970), Eaton et al. (1987) Mauna Ulu, 1969 H 300 400 0.75 Swanson et al. (1979) Miyakejima, 1983 H 185 2000 0.09 Aramaki et al. (1986) Pu’u ‘O’o, 1983–1986 H 100 315 0.3 Parfitt and Wilson (1994) Rise speeds have been calculated from observed volume fluxes and vent areas. * The eruption was described as Strombolian both on the grounds of the fall deposit it generated and the intermittent nature of the explosions. The short intervals (0.5–2 s) between explosions and the generation of fountains and an significant eruption column suggest, however, that the eruption represents a transitional event as described by Parfitt and Wilson (1995) and in the text. far through the overlying magma before the magma material that, more than anything, causes the style and itself is erupted. In effect, the gas bubbles are products of Hawaiian eruptions to differ so greatly ‘‘locked’’ to the magma in which they formed. Thus, from those of Plinian eruptions (see Parfitt and Wil- the model assumes homogeneous two-phase flow, in son, 1999). which two different fluid phases are present (the The RSD model further proposes that a different magma and gas) but in which the fluids behave as if eruption mechanism operates if the rise speed of they are a single fluid phase. In this situation, the magma is relatively low. In this case, gas bubbles growth of bubbles through diffusion and decompres- within the magma will rise upwards through the sion (Sparks, 1978; Proussevitch and Sahagian, 1996) overlying magma and can segregate from the magma and the continued formation of bubbles during ascent in which the bubbles formed (Sparks, 1978; Wilson will eventually lead to a situation in which the bubble and Head, 1981). The magma will contain a popula- volume fraction becomes large enough (f60–95%) tion of bubbles with a range of sizes—bubbles that to cause fragmentation of the magma (e.g., Sparks, formed early will have grown by diffusion and de- 1978; Wilson and Head, 1981; Houghton and Wilson, compression, while newly formed bubbles will be 1989; Thomas et al., 1994). The rising gas–magma much smaller. As the rise speed of a bubble depends mixture accelerates as it rises due to the decompres- partly on its size, a runaway situation can be achieved sion and expansion of the gas (Wilson and Head, in which an initially larger bubble, rising faster than 1981). After fragmentation, the acceleration becomes the smaller bubbles, overtakes the smaller bubbles and much more pronounced due to the reduction in wall in doing so coalesces with them. In an extreme case, friction caused by the fragmentation process and such coalescence can lead to a single large bubble that results in the eruption of a continuous jet of gas and is as wide as the conduit rising through the overlying 1 magma clots at typical speeds of f100 m sÀ (Wilson magma (essentially a slug of gas). In the RSD model, and Head, 1981). This continuous jet of material Strombolian eruptions are assumed to be the result of produces the lava fountains characteristic of Hawaiian this bubble segregation and coalescence process. eruptions (Fig. 1). As Parfitt and Wilson (1999) Wilson (1980) simulated these eruptions by consider- pointed out, this proposed mechanism is essentially ing what would happen in an open system in which the same as that envisaged as causing Plinian erup- magma was rising slowly or was static. Cooling at the tions (Wilson et al., 1980). The material erupted in top of the magma column causes the development of a Hawaiian fountains is, however, very coarse com- ‘‘skin’’ with a finite strength. The skin strength will pared with that of Plinian eruptions (Parfitt, 1998), depend on how much cooling occurs before the arrival and it is this difference in the grain size of the erupting of the large bubbles. If the interval between bubble E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 83

the character of ‘‘transitional’’ eruptions—eruptions that show aspects of both Hawaiian and Strombolian eruptions. For typical basaltic eruptions the transition between Hawaiian and Strombolian activity occurs at 1 rise speeds of f0.01–0.1 m sÀ (Parfitt and Wilson, 1995). It is expected that Hawaiian eruptions will occur at rise speeds much greater than this and Strombolian activity will occur at much lower speeds (Fig. 5).

3.2. The collapsing foam model

A series of papers published in the 1980s and Fig. 5. The controls of magma rise speed and gas content on basaltic 1990s (Vergniolle and Jaupart, 1986; Jaupart and eruption style as predicted by the RSD model. Redrawn from Parfitt Vergniolle, 1988; Jaupart and Vergniolle, 1989; Verg- and Wilson (1995). niolle and Jaupart, 1990; Vergniolle, 1996; Vergniolle and Brandeis, 1996) put forward an alternative model arrival is short enough, each bubble will updome the of the mechanisms of basaltic eruptions. The original thin skin and burst through the top of the magma paper (Vergniolle and Jaupart, 1986) challenged the column with minimal delay. If the interval between assumption of ‘‘homogeneity’’ (i.e., homogeneous the arrival of giant bubbles is longer, the skin will cool two-phase flow) made in the RSD model and pro- and thicken and then more than one bubble may have posed that both Hawaiian and Strombolian eruptions to arrive and become trapped before sufficient pres- are the result of separated, two-phase flow, i.e., sure is built up in an accumulating gas pocket to break eruptions in which the flow of the magma and gas through the skin. In either case, the short time interval phases are significantly different. They described the between explosions suggests that the strength of this different flow regimes that can prevail during sepa- skin is never very great. Repeated cycles of cooling rated, two-phase flow and proposed that Strombolian and gas accumulation followed by bubble bursting eruptions result from slug flow and Hawaiian erup- lead to the series of transient explosions characteristic tions from annular flow (Fig. 6). The model was of Strombolian eruptions. developed further by Jaupart and Vergniolle (1988) Wilson and Head (1981) presented computer mod- and Jaupart and Vergniolle (1989), wherein the con- elling to define the rise speed conditions in which ditions in which slug flow and annular flow can Strombolian and Hawaiian activity would be domi- develop were described. nant. Parfitt and Wilson (1995) carried out more In the CF model, magma is assumed to be stored detailed simulation of these conditions and discussed within some sort of storage area (a magma chamber or

Fig. 6. Schematic diagram depicting two examples of separated, two-phase flow: (a) slug flow and (b) annular flow. 84 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 a dike system) at a depth at which gas can exsolve the open conduit to become trapped by the cool ‘skin’ from the magma. The gas bubbles, once formed, rise on the top of the magma column prior to bursting. In and accumulate at the roof of the storage area and this model, there is no explicit link between magma become close-packed into a foam layer. When the rise speed and eruption style. foam layer reaches a critical thickness, it becomes Vergniolle and Jaupart (1986) challenged the as- unstable and collapses, the bubbles coalescing to form sumption made in the RSD model that coalescence of a gas pocket. The gas pocket then rises up an open bubbles can occur progressively during magma ascent. vent system and is erupted. In this model, Strombolian The RSD model assumptions are based on the obser- eruptions represent repeated partial foam collapse vation that larger bubbles rise faster than smaller ones events, whereas Hawaiian eruptions occur from com- (Fig. 7) and therefore have the opportunity to overtake plete, almost instantaneous foam collapse. In a series and coalesce with smaller bubbles. Wilson and Head of laboratory experiments, Jaupart and Vergniolle (1981) and Parfitt and Wilson (1995) assume that (1988) showed that if the viscosity of the liquid phase bubbles ‘‘which initially lie within their own radius is relatively low then the collapse of the foam is total of the vertical line of ascent of the large bubble will and the pocket of gas rises up the open conduit system make geometric contact with it’’ and will be absorbed as a single body. The observed flow is annular in this by the larger bubble. Vergniolle and Jaupart (1986) case and liquid in the annulus around the gas core is drew on work by Taitel et al. (1980) that suggests that dragged upwards with the gas and erupted (Fig. 6). coalescence only occurs when bubbles are rising fast Jaupart and Vergniolle (1988) liken this behaviour to enough to deform during ascent. This work suggested that of a Hawaiian eruption. If the viscosity of the that only bubbles larger than f40 mm will be able to liquid is higher, the foam collapses only partially and coalesce with smaller bubbles. As bubbles only reach forms a series of smaller gas pockets. These travel up sizes of 10–50 mm by decompression and diffusion the conduit system periodically in slug flows and Vergniolle and Jaupart (1986) argue that bubble coa- burst at the surface. This behaviour is likened to lescence cannot occur during ascent, i.e., that the RSD Strombolian eruptions. model is invalid. More recent work by Manga and Stone (1994), however, suggests that bubbles >5 mm radius will deform during ascent and that such bubbles 4. Strombolian eruptions enhance coalescence, i.e., that coalescence can occur with bubbles of smaller size but that if larger bubbles The RSD and CF models do not differ very much are present, models such as that of Wilson and Head in their view of Strombolian activity. They both treat (1981) will underestimate the amount of coalescence these eruptions as occurring when gas segregates from that occurs. So coalescence can occur for smaller the magma and accumulates as a gas pocket that then bubbles, but once bubbles have grown to sizes >5 bursts at the top of an open magma column producing mm enhanced coalescence will facilitate runaway the mild explosions characteristic of Strombolian coalescence. Evidence from the study of bubble size activity. This behaviour is consistent with direct distributions in lava and tephra supports the idea that observations of eruptions (e.g., Vergniolle and Bran- bubble coalescence occurs during magma ascent (e.g., deis, 1994) and studies of the acoustic wave that Mangan et al., 1993). accompanies each explosion (Vergniolle and Bran- While the assumptions in Wilson and Head (1981) deis, 1994, 1996; Vergniolle et al., 1996). and Parfitt and Wilson (1995) about whether two The main difference between the models concerns bubbles will coalesce are obviously a simplification where gas accumulation occurs within the magmatic of the real situation, the current evidence does suggest system. In the RSD model, the gas segregation is that it is possible for coalescence to occur in rising considered to be progressive, with bubble coalescence magmas as long as the bubbles have the opportunity to occurring because the magma rise speed is low. By move upwards relative to the magma, i.e., as long as contrast, in the CF model bubbles are assumed to the magma rise speed is low. A link between explosive accumulate at some depth forming a foam layer that basaltic eruption style and rise speed is evident from then partially collapses (or coalesces) and travels up field observations of recent eruptions (e.g., Table 1). E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 85

2 Fig. 7. The relationship between bubble radius and bubble rise speed through magma. The rise speed, u, was calculated from u=(2(qm –qg)gr )/ 9g where qm and qg are the magma and gas densities, g is the acceleration due to gravity and g is the magma viscosity. Line 1 represents the case 3 in which the magma is assumed to have a density of 2600 kg mÀ and a viscosity of 10 Pa s. Line 2 represents the case in which the magma is 3 assumed to have a density of 2000 kg mÀ and a viscosity of 30 Pa s.

Furthermore, Parfitt and Wilson (1995) suggested that seismic signal that consists of an initial compression for typical magma volatile contents the transition from followed by a dilation and further compression (Neu- Strombolian to Hawaiian activity occurs between rise berg et al., 1994; Chouet et al., 1999). Chouet et al. 1 speeds of 0.01 and 0.1 m sÀ (Fig. 5). Comparison (1999) have shown that the seismic source varies in with the examples given in Table 1 show that (a) depth through the course of the explosion, starting at a Strombolian eruptions are indeed associated with low- depth of 125 m, deepening to a depth of f350 m and er rise speeds and (b) that the transition in eruption then shallowing again to a depth of around f200 m. style occurs within the rise speed range predicted by They suggest that this seismic event is caused by the Parfitt and Wilson (1995). This would seem to support uprush of a gas pocket of the sort pictured in the CF their contention that coalescence is progressive and model, though it has yet to be demonstrated that the dependent on the magma rise speed. details of the seismic signal are consistent with the In contrast to the RSD model, the CF model of upward passage of a gas slug. There is therefore no Strombolian eruptions requires that gas segregation definitive answer at present as to which gas segrega- and foam formation occurs during storage at depth tion process operates at Stromboli. In a broader and thus can only operate under the particular circum- context, there is no reason why one model should stance where a storage zone exists beneath the vent at explain all Strombolian activity. It must be borne in a depth at which exsolution of one or more gas phases mind, however, that the CF model can only apply in a is occurring. At Stromboli itself, there is evidence that particular combination of circumstances—where there magma storage can occur at depths no greater than a is a suitable storage zone at a depth where one of more few hundred metres (Giberti et al., 1992). Each gas phase can exsolve—whereas the RSD model is explosion at Stromboli is associated with a distinct applicable to any open system. 86 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107

Both models assume that shallow bubble bursting 5. Hawaiian eruptions causes the observed explosions so that the same model of the bursting process is applicable in either The RSD and CF models present very different case. Wilson (1980) modelled the ejection of clasts in views of the dynamics of Hawaiian activity. Two fun- Strombolian eruptions by assuming that the eruptions damental differences exist between the models. These result from the bursting of near-surface bubbles. The are concerned with the nature of the fluid flow at depth model links the initial pressure in the bursting bubble, and with the dominant volatile species driving the the weight percentage of gas erupted and the maxi- eruptions. Each difference is considered here in turn. mum velocity achieved by the ejected matter (Fig. 8). Pressures within the bursting bubbles are unlikely to 5.1. Flow regimes in Hawaiian eruptions exceed 0.3 MPa (Blackburn, 1977; Sparks, 1978). Observations at Heimaey and Stromboli suggest that The RSD model assumes that homogeneous two- the weight percentage of gas in typical explosions is phase flow prevails. By contrast, the CF model 10 –30 wt.% (Blackburn et al., 1976), although assumes that separated two-phase flow occurs. There Chouet et al. (1974) note that some events at Strom- are a range of flow regimes in which separated two- boli can have gas contents as high as 94 wt.%. Direct phase flow can occur, and the CF model assumes that observations suggest that clasts are ejected in some annular flow (Fig. 6) prevails during Hawaiian erup- 1 Strombolian eruptions at speeds of up to 230 m sÀ tions (Vergniolle and Jaupart, 1986). I now discuss the (Blackburn et al., 1976). At Stromboli itself, speeds implications of, and evidence for, each type of flow. 1 are more typically <100 m sÀ (Chouet et al., 1974; The assumption of homogeneous two-phase flow is Blackburn et al., 1976). Comparison of these values never strictly valid because gas bubbles are always with the model results in Fig. 8 shows that there is buoyant relative to the magma and thus are always broad consistency between the model predictions and rising faster than the magma. However, as stated field observations. above, if the rise speed of the magma is rapid the bubbles do not rise far through the overlying magma before the magma is erupted and in effect the gas bubbles are ‘‘locked’’ to the magma, i.e., the assump- tion of homogeneous flow is valid. Thus, it is the rise speed of the bubbles relative to the magma rise speed which determines whether flow is homogeneous or not. The rise speed of a bubble depends on its size, larger bubbles rising faster than smaller ones (Fig. 7). The validity of the assumption of homogeneity depends, therefore, on the size of the bubbles involved and on the magma rise speed at depth. Table 1 shows that rise speeds in Hawaiian eruptions are typically 1 >0.1 m sÀ . Fig. 7 shows that only at radii of z0.01 m (10 mm) does the bubble rise speed through the magma become of the same order of magnitude as the magma rise speed. For bubbles with radii <5 mm, the bubble rise speed is always likely to be more than an order of magnitude less than the typical magma rise speed in a Hawaiian eruption. Thus, the assumption of Fig. 8. Diagram showing the relationship between bubble pressure homogeneity is likely to be valid as long as the bubble and the maximum ejecta velocity in a Strombolian eruption. The radii are less than f5 mm. So, the crucial issue is the different curves represent different weight percentages of gas in the erupted material. The cross-hatched area represents the likely range size of the bubbles within the rising magma. of conditions during Strombolian eruptions. Redrawn from Wilson Bubbles form in magma when the magma becomes (1980). supersaturated in the volatile concerned. The depth E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 87 beneath the surface at which bubbles start to form (1996) report radii for bubbles in basaltic scoria from depends on the amount of dissolved volatiles and the the Pu’u ‘O’o eruption of Kilauea of V2.5 mm. While species of volatile involved (Wilson and Head, 1981). we cannot know for sure how such sizes relate to Bubbles have typical radii of f10 Am when first bubble sizes prior to fragmentation, certainly such formed and then grow by diffusion and decompression studies do not provide any compelling reason to think as the magma rises (e.g., Sparks, 1978). Sparks (1978) that bubbles sizes in basaltic magmas exceed the f5 presented numerical modelling of the growth of bub- mm size predicted by the theoretical studies. bles by diffusion and decompression and found that Bearing all these points in mind it seems reason- maximum bubble sizes for H2O bubbles exsolving able to assume that the radii of the majority of H2O from a basaltic magma depend on the amount of bubbles in magma with a water content V0.5 wt.% is dissolved water in the magma. For a gas content of likely to be b5 mm. This means that in the case of 0.5 wt.% (a reasonable value for basaltic magmas), the H2O bubbles in basaltic magma, the situation consid- maximum radius is f5 mm. A more recent study by ered in all the RSD modelling, the assumption of Proussevitch and Sahagian (1996) gives maximum homogeneity is almost certainly valid. bubble radii of 6–8 mm for H2O bubbles in basaltic The RSD modelling has only considered the situ- magma using initial water contents of 1.52% and 3.03 ation of water exsolution. It is important to note, wt.%, respectively. The sizes would be smaller for though, that the situation would be different in the more reasonable initial water contents. Thus, theoret- case of CO2 exsolution. CO2 is less soluble in magma ical studies suggest that water bubbles forming in than water and so exsolves and forms bubbles at rising basaltic magmas would typically have maxi- greater depths beneath the surface. This means that mum radii of f5 mm for water contents typical of CO2 bubbles experience more growth by decompres- most basaltic eruptions. The size of the largest bubble sion during ascent than do H2O bubbles. Fig. 9 shows is not, however, representative of the bubble popula- that for CO2 contents in the range of 0.1–0.5 wt.% tion as a whole. Most bubbles will reach an interme- (reasonable values for a basaltic magma), bubbles are diate size. For instance, Sparks (1978) showed that for bubbles formed in basalt containing 1 wt.% water the maximum radii would be f40 mm but the typical size would be 1–10 mm rather than 40 mm. Furthermore, in the modelling studies just described, it is assumed that bubbles continue growing all the way to the surface. In practice, though, magma fragmentation will occur beneath the surface and so the maximum bubble size will not be achieved. These theoretical studies suggest then that for typical water contents the typical size of bubbles in basaltic magmas will be b5 mm. Determining the bubble sizes in real magmas is extremely problematic because the fragmentation pro- cess destroys much of the evidence of pre-fragmenta- tion bubble sizes. A number of studies have looked, though, at sizes of bubbles in basaltic scoria and lava (e.g., Cashman and Mangan, 1995; Mangan and Cashman, 1996). Bubbles contained in such samples represent bubbles formed in magma clots after frag- mentation but also bubbles which survived the frag- Fig. 9. Diagram showing the relationship between final bubble size and magma rise speed for magma containing 0.1, 0.3 and 0.5 wt.% mentation process and continued to grow after 1 CO2. At rise speeds less than f1 m sÀ , the bubbles are able to rise fragmentation. Cashman and Mangan (1995) report through the overlying magma and in doing so to coalesce. The mean bubble radii for quenched lava from Kilauea resulting bubbles are considerably larger than those developed in volcano of 0.1–0.15 mm and Mangan and Cashman faster rising magma where bubble coalescence is negligible. 88 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 likely to experience coalescence at rise speeds of V1 erupted per second is 2.6 105 kg (assuming a magma 1 3 msÀ (and so achieve final diameters anywhere in the density of 2600 kg mÀ ) and so the mass of water range 3 mm to 10 m). So, as coalescence is evidence released from this magma during ascent is 1300 kg. At for separated flow, and as rise speeds typical of atmospheric pressure this mass of gas occupies 8667 1 3 basaltic eruptions are generally <1 m sÀ (Table 1), m (the density of steam at atmospheric pressure and 3 we would expect separated two-phase flow to occur magmatic temperature is f0.15 kg mÀ ). So, at the during ascent. In the case of CO2 exsolution, then, it surface, the volume of the magma compared with the would be invalid to assume that homogeneous two- volume of gas is f 1%, even though there has been phase flow occurs. As I have stated, however, CO2 no concentration and segregation of the gas from the has not been considered as the ‘driving’ gas in any of magma prior to eruption. It is the mass of magma the RSD modelling. The issue of which volatile acts relative to the mass of gas erupted that is crucial as the ‘driving’ gas for Hawaiian eruptions is dis- evidence of segregation or homogeneity, not the cussed in detail below. volume. Vergniolle and Jaupart (1986) argued that homo- This point can be further tested using a real geneous two-phase flow does not occur during Ha- example. Between 1983 and 1986, a series of 47 lava waiian eruptions. They based this assertion on a fountaining episodes occurred at Pu’u ‘O’o, a vent on number of lines of evidence. The first is that the the flanks of Kilauea Volcano (Heliker and Wright, characteristic radius of bubbles in Hawaiian eruptions 1991). During a number of episodes, measurements is 50 mm. Such a bubble would have a rise speed were made of the mass of CO2 and SO2 released and 1 through the magma of f1 m sÀ (Fig. 7) and thus a of the relative volumes of each gas species released in speed that is comparable to the magma rise speed at each eruption (e.g., Greenland et al., 1985; Greenland, depth (Table 1). In such a situation, the bubbles would 1988). By combining these measurements, it is pos- tend to separate from the magma and the assumption sible to estimate the mass of each gas species released of homogeneous two-phase flow would break down. during each episode. As measurements were also As explained above, such bubble sizes are only likely made of the volumes of lava erupted in each episode to be achieved in eruptions in which CO2 is the (e.g., Wolfe et al., 1988), it is possible to assess driving gas. Thus, as just stated, the crucial issue is: whether the amounts of gas released are in excess of Which gas species ‘drives’ these eruptions? This is that originally dissolved in the magma: If the CF discussed in more detail below but the initial conclu- model is valid, the gas mass fraction in the erupted sion that can be drawn is that homogeneous flow is material will be considerably greater than that in the possible in Hawaiian eruptions driven by H2O but not magma at depth. Consider, then, one example from those driven by CO2. this eruption. Episode 16 of the eruption (in March This bubble size argument is not the only one 1984) produced H2O at a rate of 40000 tonnes/day presented by Vergniolle and Jaupart (1986) to support and CO2 at 3200 tonnes/day (Greenland et al., 1985). their contention that separated rather than homoge- The eruption lasted for 31 h, so a total of 51700 7 neous two-phase flow occurs during Hawaiian erup- tonnes (5.17 10 kg) of H2O and 4130 tonnes 6  tions. Another argument concerns the volumes of gas (4.13 10 kg) of CO2 were released during this  6 and magma present upon eruption. They note that episode. The volume of lava produced was 12 10 3  magma typically makes up less than f1% of the m (Wolfe et al., 1987), which, assuming a lava bulk 3 erupted volume in a Hawaiian lava fountain and argue density of f2000 kg mÀ , yields an erupted mass of that such a situation cannot be achieved in an eruption magma of 2.4 1010 kg. This yields gas mass fractions  in which homogeneous two-phase flow prevails. This in the erupted material of 0.22 wt.% of H2O and 0.017 argument is fundamentally flawed, as can be demon- wt.% of CO2. Residual gas contents in Kilauean lavas strated by the following simple calculations. are typically 0.10 wt.% H2O and 0.015 wt.% CO2 for Consider a basaltic magma exsolving 0.5 wt.% Kilauea (Gerlach and Graeber, 1985), which yields water during ascent. In an eruption with a magma estimates of the gas content within the magma prior to 3 1 volume flux of 100 m sÀ (the situation treated by eruption of 0.33 wt.% H2O and 0.032 wt.% CO2. Vergniolle and Jaupart, 1986) the mass of magma Similar calculations for other Pu’u ‘O’o episodes E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 89 produce similar results. Independent estimates of ‘drives’ Hawaiian eruptions. The RSD model as- volatile contents based on fluid inclusions studies sumes that the ‘driving’ gas is H2O, whereas the give H2O contents for the Pu’u ‘O’o eruption of CF model assumes that it is CO2. This issue is crucial 0.39–0.51 wt.% for tephra from the high fountain because, in the situations considered in the published events and 0.10–0.28 wt.% for spatter from less models, the RSD model is incompatible with the vigorous activity (Wallace and Anderson, 1998). So driving gas being CO2 and the CF model is incom- the gas released during the eruptions is consistent with patible with the driving gas being H2O. It is possible, the gas contents contained within the magma prior to therefore: eruption, and, thus, there is no evidence to support the (a) to distinguish between the two models for idea that gas concentration and separation occurred specific eruptions provided observational evidence prior to eruption. The values instead support the exists about the species and mass fractions of gas contention of the RSD model that Hawaiian eruptions released in the eruption (see below); and result from homogeneous two-phase flow. It is also (b) that each mechanism could be valid in different worth noting that in this eruption, the volume per- volcanic situations depending on the gas species, mass centage of the magma in the lava fountain is f0.35% fraction present in the magma, and the storage history (calculated in the same way as in the example given of the magma as it ascends. above). This supports my contention that, even in a Let us examine why the two models assume homogeneous eruption, the volume percent of magma different ‘driving’ gas species and then look at which in the fountain can be <1%, and, thus, that the situation is most common in actual eruptions. statement by Vergniolle and Jaupart (1986) that this is evidence of separated flow is erroneous. 5.2.1. H2O as the ‘driving’ gas More fundamentally, Vergniolle and Mangan (2000) The RSD model assumes in all cases that the gas describe a distinctive pattern of behaviour observed driving Hawaiian eruptions is H2O. This is for two during the 1959 Kilauea Iki eruption in which magma main reasons: was simultaneously erupted in a lava fountain and (1) Water is usually the most abundant volatile drains back around the edges of the vent. They assert present within basaltic magmas (e.g., Wallace and that this observation is evidence for annular flow and Anderson, 2000). that simultaneous drainback and eruption is not possi- (2) Water only exsolves from basalts at shallow ble during homogeneous flow. Wilson et al. (1995) depths (typically a few hundred metres) beneath the have previously published a model in which simulta- surface (Sparks, 1978; Wilson and Head, 1981 ). This neous drainback and eruption occurs during homoge- means that the water will usually have had little neous flow. This issue has been examined again by opportunity to exsolve and escape from the magma Lionel Wilson (unpublished calculations, 2003) and his as it ascends towards the surface, and thus its exso- findings are contained in Appendix A. His treatment lution from the magma near the surface must play shows that it is perfectly possible to explain the some role in the eruption dynamics. observation of simultaneous drainback and eruption at Kilauea Iki in terms of homogeneous flow. 5.2.2. CO2 as the ‘driving’ gas In conclusion, arguments presented as evidence The CF model assumes that the driving gas is CO2. that separated two-phase flow must occur during This is because in this model gas bubbles must Hawaiian eruptions do not stand up to detailed scru- accumulate as a foam layer in a storage area at depth tiny. Existing observational evidence, instead, sup- in order for separated two-phase flow to occur. For ports the contention that Hawaiian eruptions occur this to be possible, it is necessary that: as the result of homogeneous two-phase flow. (1) Storage occurs at a depth where exsolution of the ‘driving’ gas can occur. 5.2. Dominant volatile species (2) The roof of the storage zone has sufficient area to allow the accumulation of a sufficient volume The other fundamental difference between the two of foam to be consistent with observed erupted models concerns the species of volatile that typically volumes. 90 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107

The shallow exsolution depths of H2O make it accounted for only 3% of the volatiles released extremely unlikely that these criteria will be met, (Greenland, 1984; 1988). It is difficult to accept, whereas CO2 exsolves at depths of several kilometres therefore, that CO2 is the ‘driving’ gas rather than beneath the surface and thus within zones where H2O. large-scale storage often occurs. Although it is possible for CO2 to be exsolved and 5.3. Initial conclusions stored in the way the CF model suggests, Parfitt and Wilson (1994) have pointed out that a problem with The RSD and CF models of Hawaiian eruptions this model is that it neglects the effects of water make fundamentally different assumptions about the exsolution, occurring as the foam ascends, will have flow regime prevailing at depth and about the on eruption. Their argument is that, even if the volatiles driving the eruptions. Both models could eruption were driven from depth by foam collapse, potentially apply in different situations depending on the magma that is carried up and erupted will still the volatiles species, bubble sizes, storage history contain dissolved water (water being abundant in and magma rise speeds concerned. Neither model basalts). This water will exsolve from the magma as appears to have any fundamental flaw. However, the it reaches shallow levels and thus this water must play usefulness of a model depends not on its theoretical some role in driving the eruption. Vergniolle and validity but on how well it reproduces the activity Jaupart (1986) have argued that, though the water which occurs in nature, and in this respect, observa- exsolves from the magma as it rises, ‘‘the small tional data examined thus far favour the RSD model vesicles formed through exsolution in the conduit over the CF model. cannot coalesce and can therefore reach high volume Both models have been used to look at the same fractions without leading to a change in flow regime’’. test case—the 1983–1986 Pu’u ‘O’o eruption—and In other words, water bubbles do form in the magma both models purport to explain the observational data as it rises towards the surface but this exsolution is collected during that eruption. As the two models ‘‘passive’’ because it generates magma clots with high make fundamentally different assumptions and pre- vesicularity, but this material does not fragment or in dictions about Hawaiian activity, it is impossible that any way drive the eruption. This explanation appears both models are consistent with the same set of to be flawed in two ways: observational data. For this reason, I will now exam- (1) In the example given above, it was shown that ine, in detail, how the models have been tested using the volume of the water exsolved from the rising evidence from this eruption. magma compared with the volume of the magma from which it exsolved is such that at the vent the magma represents V1% of the total volume at the vent. If the 6. The 1983–1986 Pu’u ‘O’o eruption of Kilauea exsolving water is held, as Vergniolle and Jaupart Volcano (1986) suggest, as small bubbles within the magma this means that the vesicularity of the erupting magma This eruption started in January 1983 with the would have to exceed 99% in all of the erupted emplacement of a feeder dike laterally from the magma. The most vesicular material generated in summit magma chamber into Kilauea’s East Rift Hawaiian eruptions, reticulite, has vesicularities rang- Zone (ERZ) (Klein et al., 1987; Wolfe et al., ing up to 98% (Thomas et al., 1994) but reticulite 1987). The dike fed a fissure eruption on the middle makes up only a small proportion of the material ERZ at distances of 14–22 km from the summit. produced in Hawaiian eruptions. Dike emplacement and eruption were accompanied (2) If the water is trapped in small bubbles within by major deflation of the summit magma chamber the clasts, then it would not be released in the eruption (Fig. 10). After about a month, during which the plume. Yet, in the Pu’u ‘O’o eruption, which Verg- summit magma chamber reinflated (Fig. 10), a new niolle and Jaupart (1990) and Vergniolle (1996) use as eruptive episode began in the same area of the ERZ a test case for the CF model, 85% by volume of the fed through the same feeder dike (Wolfe et al., measured volatile release was water whereas CO2 1987). A pattern of activity developed in which E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 91

Fig. 10. The summit tiltmeter record for Kilauea volcano for 1983. The numbers indicate each eruptive episode of the Pu’u ‘O’o eruption during 1983. Redrawn from Wolfe et al. (1987).

eruptions typically f1 day in duration occurred istics of the eruption that both models seek to associated with deflation of the summit and punctu- explain. ated by repose periods of f3 weeks during which the summit reinflated. By the 4th eruptive outbreak, 6.1. The cyclic character of the eruption activity had become concentrated at one eruptive vent subsequently named Pu’u ‘O’o (Wolfe et al., A key feature of the eruption was its repetitive, 1988). The eruption continued this cyclic pattern cyclic character. Each eruption was preceded by a until July 1986 when the location and behaviour of repose period during which slow inflation of the activity switched to a vent 3 km further down rift summit occurred accompanied by minor explosive (Heliker and Wright, 1991). Eruption at this new activity at the vent and each eruption was accompa- vent, Kupaianaha, was characterised by continuous nied by rapid deflation of the summit in association minor explosive activity and slow outpouring of with high lava fountaining and generation of lava lava. The eruption was monitored in great detail by flows (Fig. 10). Dvorak and Okamura (1985) ob- the staff of the Hawaiian Volcano Observatory served that the deflation rate increased as the eruption (HVO) and their observations have been published sequence continued while the duration of each erup- in a number of papers (e.g., Dvorak and Okamura, tive episode gradually decreased (Fig. 11). They 1985; Wolfe et al., 1987; Greenland, 1988; Okamura suggested that this behaviour reflected an evolution et al., 1988; Wolfe et al., 1988; Heliker and Wright, of the magma system feeding the eruption. Parfitt and 1991; Heliker et al., 2003). Thus, this is an eruption Wilson (1994) noted that the deflation during each for which there is an exceptionally large and com- episode showed a characteristic pattern in which the plete set of field and geophysical observations with rate was initially low, increased to a peak value, and which to test the eruption models. then declined approximately exponentially (Fig. 12). Vergniolle and Jaupart (1990) and Vergniolle Parfitt and Wilson (1994) adopted the interpretation (1996), and Parfitt and Wilson (1994), have pre- that inflation and deflation of Kilauea’s summit mag- sented very different, and mutually incompatible, ma chamber occurs primarily as the result of the inflow models of this eruption based on the RSD and CF and outflow of magma (e.g., Dzurisin et al., 1984; models described above. I now compare the two Dvorak and Dzurisin, 1993). The idea is that magma is models by looking at some of the key character- supplied to the magma chamber from deeper levels at a 92 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107

Fig. 11. (a) The maximum deflation rate and (b) the duration of each episode during the 1983–86 Pu’u ‘O’o eruption. Data courtesy of the Hawaiian Volcano Observatory.

3 1 fairly constant rate (estimated at f3m sÀ ; Dzurisin various geometries. By assuming that the dike was of et al., 1984; Dvorak and Dzurisin, 1993). This leads to non-uniform geometry they were able to reproduce the slow chamber inflation during times when no high observed deflation patterns (Fig. 12), to explain why fountaining was occurring. When an eruption occurs, the eruptive behaviour was cyclic and to examine the magma is withdrawn and erupted at a rate that exceeds factors which determined when each episode started the inflow rate from the mantle and thus rapid deflation and stopped. In their model, the cyclic nature of the occurs. Starting with this premise, Parfitt and Wilson eruption is determined by the details of the subsurface (1994) examined the deflation patterns that would storage and movement of magma not by the eruption result from flow of magma through feeder dikes of style (as is the case in the CF model—see below). E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 93

common in basaltic eruptions where the initial high pressure in the chamber allows high flow rates near the start of an eruption, flow rate gradually declining as the chamber pressure declines (Wadge, 1981). In the case of the Pu’u ‘O’o high fountaining episodes, the flow rate through the dike system, and hence the rise speed beneath the vent, is sufficiently high to allow homo- geneous two-phase flow. During the repose periods between high fountain episodes, the flow rate through the dike system becomes negligible and so the rise speed beneath the vent is close to zero. In this situation, gas segregates from the magma within the vent and rises to the surface giving rise to the minor explosive activity which characterised the repose peri- ods (Wolfe et al., 1987, 1988). Vergniolle (1996) interpreted the cyclic pattern of the eruption and the associated inflation/deflation patterns in a very different way. In her model at least part of the inflation and deflation is viewed as result- ing from changes in gas volume in the summit magma chamber. Inflation is related to exsolution of CO2 from the stored magma and its accumulation as a foam layer at the roof of the magma chamber. Col- lapse of this foam layer triggers eruption of magma and deflation of the magma chamber. There are a number of problems with this model: (1) As mentioned above, observation shows that Fig. 12. Patterns of summit deflation during the Pu’u ‘O’o eruption. CO2 constitutes an average of only f3% of the total During each episode the summit deflated at a rate which was volume of gas released in the eruptive episodes initially slow, increased rapidly to a maximum value and then (Greenland, 1984; 1988). The majority of the gas declined approximately exponentially until the eruptive episode released is magmatic water (85%), which must play ended. The patterns of deflation during two eruptive episodes are a significant role in the eruption but cannot be shown—episodes 11 and 31. The bold line shows the actual deflation rate derived from the summit tiltmeter records kept by the collected as a foam prior to these eruptive episodes Hawaiian Volcano Observatory. The dashed lines represent the (see above). modelled deflation rate calculated for each episode using a model (2) The Vergniolle (1996) model requires that CO2 developed by Parfitt and Wilson (1994). The diagram is modified exsolving within the magma chamber should become from Parfitt and Wilson (1994). trapped at the chamber roof forming the foam layer that ultimately causes each fountaining episode. Ger- lach and Graeber (1985), Gerlach (1986) and Gerlach Instead, the Hawaiian character of the eruption is and Taylor (1990) have studied gas release from the controlled by the cyclicity because the cycles are Kilauea system and show that magmas erupted on the related to variations in flow rate through the dike rift zones, including the Pu’u ‘O’o magma, are system and thus to the rise speed of the magma beneath depleted of CO2 prior to eruption (consistent with the vent. The flow rate through the dike system is Point 1). They propose that CO2 is lost from the directly correlated with the deflation rate; thus the flow magma chamber during storage and show that the 6 rate rapidly increases as cooled magma is pushed measured daily release rates of CO2 (1.6 to 3.6 10 1  through the dike system, reaches a peak and then kg dayÀ —Greenland et al., 1985) from the summit declines exponentially (Fig. 12). Such a pattern is region are consistent with calculated release rates that 94 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 are based on the influx of magma into the chamber of the eruption to 0.6 109 m3 for the last high 6 1  over a period of decades (f3.7 10 kg dayÀ — fountaining episode. Such values are almost an order  Gerlach and Graeber, 1985). Gerlach and Graeber of magnitude greater than volumes calculated from (1985), Gerlach (1986) and Gerlach and Taylor measurements of gas mass release during the eruptive (1990) conclude that most of the magmatic CO2 is episodes (Table 2). Thus, the data presented in Verg- lost from the magma chamber through non-eruptive niolle and Jaupart (1990) as evidence for a decline in degassing. So, although CO2 does exsolve in the gas release through the eruption sequence must be magma chamber, the observational evidence suggests treated with scepticism. Furthermore, observational that it is lost from the chamber by degassing rather evidence does not support the idea that less gas was than being trapped at the roof of the chamber to being released during the continuous phase of activity produce a foam layer as the CF model requires. at the Kupaianaha vent compared with the high foun- (3) The model of Vergniolle and Jaupart (1990) and taining phases at Pu’u ‘O’o which preceded it. Meas- Vergniolle (1996) explains the change in eruptive urements of SO2 emission rates during the Kupaianaha behaviour in July 1986 from cyclic fountaining to eruption show that the rates are 5–27 times less than continuous lava outpouring as being the result of a the emission rates during the high fountaining episodes decline in gas volume through time. Values for the (Andres et al., 1989). However, eruption rates at declining gas release in each episode are shown in Fig. Kupaianaha are also lower, averaging 0.35 106 m3 1 6 3 1  7 in Vergniolle and Jaupart (1990). These data were dayÀ compared with f7.7 10 m dayÀ during the  derived by Vergniolle and Jaupart (1990) from meas- high fountaining episodes (Heliker et al., 2003), i.e., urements of the maximum fountain height recorded for 22 times less than during high fountaining. Thus, the each episode and the total eruption duration made by decreases in emission rates and eruption rates are HVO staff. The gas volume was calculated by obtain- comparable. Averaged over time the continuous slow ing the exit velocity for each episode from the fountain release of gas and magma from Kupaianaha actually height and then multiplying this by the vent cross- released as much gas as the higher rate but short-lived sectional area and the eruption duration. There are high fountaining episodes. Thus, there is no evidence several reasons why this is inaccurate way of estimat- ing the gas volume released: (a) The maximum fountain height is not repre- sentative of the episode as a whole and represents a Table 2 time when the exit velocity is a maximum. This is Gas volumes released during episodes 15 and 16 of the Pu’u ‘O’o eruption clear from time-lapse data collected by HVO, some Gas Gas mass Total gas Gas density Gas volume of which was published in Wolfe et al. (1987, 3 3 species released mass released (kg mÀ ) erupted (m ) 1988). per day during the (b) The estimates of large gas volumes during the (tonnes/day) episode (kg) early episodes of the sequence, which add greatly (a) Gas release during episode 15. The episode duration was 19 h to the impression that gas volume declines through H O 58,000 4.59 107 0.15 3.06 108 2   SO 27,000 2.14 107 0.53 4.03 107 time, result from using the long durations of these 2   CO 4700 3.72 106 0.36 1.03 107 episodes to calculate the volumes. Observational 2   HCl 330 2.61 105 0.3 8.71 105 evidence shows, however, that the vents were not   HF 200 1.58 105 0.16 9.90 105   active throughout the duration of the episode and Total 3.59 108 thus the use of the total durations to calculate the  volumes is inappropriate. (b) Gas release during episode 16. The episode duration was 31 h H O 40,000 5.17 107 0.15 3.44 108 (c) Finally, the calculation takes no account of 2   SO 18,000 2.33 107 0.53 4.39 107 the expansion of the gas as it rises. 2   CO 3200 4.13 106 0.36 1.15 107 2   That the values of gas volume calculated by Vergniolle HCl 220 2.84 105 0.3 9.47 105   and Jaupart (1990) are unreliable can be verified by HF 140 1.81 105 0.16 1.13 106   comparison with observational data. Their gas vol- Total 4.02 108 9 3  umes range from 3.1 10 m during the initial stages Gas masses released are taken from Greenland et al. (1985).  E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 95 that any less SO2 was being released from the magma and style of eruption changed abruptly. A fissure after the change in eruption style. Though no evidence system opened up that extended downrift from Pu’u is available for release of other volatiles before and ‘O’o and activity from this system eventually local- after the change in eruptive character, it seems unlikely ised at a new vent later named Kupaianaha (Heliker that the overall release rate should decline while the and Wright, 1991). The change in locality corre- SO2 release rate remains unchanged. sponded to the end of the cyclic lava fountaining activity seen at Pu’u ‘O’o. Instead, the activity be- 6.2. Fountain heights and exit velocities came continuous and occurred at a much slower 3 1 eruption rate (f5m sÀ ). A low lava shield with a Measurements were made by HVO staff of the lava lake at the top gradually developed. Lava within maximum and average fountain heights for each the lake circulated and degassed (Fig. 13a) and was eruptive episode (Wolfe et al., 1987, 1988). Wilson continually drained from the lake through a complex and Head (1981), Head and Wilson (1987) and Parfitt tube system (Mattox et al., 1993). Though some et al. (1995) have related fountain heights to the degassing occurred at Kupaianaha, the bulk of the eruption rate and gas content of the erupting magma gas release occurred through the Pu’u ‘O’o vent as using the RSD model by calculating the exit velocity was evident from observation of a plume constantly of the erupting mixture and assuming that the larger rising from the cone (Fig. 13b) and confirmed by clasts (which form the main part of the fountain) direct measurements (Andres et al., 1989). This behave ballistically. When this model is applied to change in character is similar to ones which occurred the Pu’u ‘O’o eruption, it suggests that the observed during the 1969–1974 Mauna Ulu eruption (Swanson fountain heights would be produced if the water et al., 1979; Tilling et al., 1987). Vergniolle and content of the erupting magma is 0.32 wt.%. This is Jaupart (1990) and Vergniolle (1996) propose that consistent with independent estimates that range from changes from high fountaining to continuous eruption 0.21% to 0.38 wt.% (Gerlach and Graeber, 1985; in each of these eruptions represent a change in Greenland et al., 1985; Greenland, 1988). eruption style from Hawaiian to effusive. Parfitt and The CF model (Vergniolle and Jaupart, 1990; Wilson (1994) have argued that the change in eruption Vergniolle, 1996) does not make a prediction of the character represents a change from Hawaiian to exit velocities or fountain heights of the eruption. Strombolian. Thus, there is a basic disagreement about how to interpret the observed activity as well 6.3. Volumes and durations as a disagreement on the causes of the change. Part of the problem arises because of the unusual nature of Observational evidence collected by HVO staff gas release during the Kupaianaha eruption. The (Heliker and Mattox, 2003) provides constraints on magmatic plumbing system established between Pu’u the volumes of lava produced during each episode (2 ‘O’o and Kupaianaha in July 1986 allowed shallow to 38 106 m3), on the average eruption rates (12 to degassing of the magma through the Pu’u ‘O’o vent Â3 1 489 m sÀ ) and on the duration of each episode (5 to prior to magma eruption at Kupaianaha. This means 290 h). Parfitt and Wilson (1994) used the RSD model that although there was minor explosive activity at to simulate the Pu’u ‘O’o episodes and the model can Kupaianaha (Fig. 13a), the overflow of the lake can be adequately explain the observed values of each of interpreted as effusive activity. Parfitt and Wilson these parameters. It has never been demonstrated that (1994) argue, however, that the eruption is Strombo- the CF model can explain these observed eruption lian because the observation of significant gas release volumes or durations. and spattering within the Pu’u ‘O’o cone (Andres et al., 1989; Mangan et al., 1995) shows that gas is 6.4. Change in eruption character segregating in significant quantities and giving rise to explosive activity at the top of the magma column. A further fundamental difference between the two Effusive activity corresponds to events where no models is highlighted by the change in eruption significant gas segregation is occurring. This is more character that occurred in July 1986, when the site evident when considering the Mauna Ulu eruption. 96 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107

Fig. 13. (a) Photograph of the Kupaianaha lava lake. The lava lake is covered with a cooled crust which was constantly moving and overturning. A large bubble in the process of bursting can be seen near the far wall of the lake. Gas release from the lake was constant and sufficient to prevent an observer watching the activity for periods of more than a few minutes at a time. (b) The Pu’u ‘O’o cone viewed from Kupaianaha. The photograph was taken at the same time as that in (a). It is evident from the plume rising from Pu’u ‘O’o that significant quantities of gas were being released there while eruption occurred from Kupaianaha. (Both photographs taken by the author, February 20, 1988).

Here, after the change from high fountaining to the magma chamber as the magma became depleted of continuous activity, eruption was associated dome gas. As we have seen above, the evidence that the gas fountaining, gas-pistoning, spattering and low foun- release rate is smaller after the change in eruption taining, all of which indicate that gas release and character is unconvincing. Furthermore, this argument minor explosive activity was associated with the is based on the idea that the magma chamber is not production of lava (Swanson et al., 1979). being resupplied with magma. Many studies suggest Vergniolle and Jaupart (1990) and Vergniolle that the magma chamber is fairly continuously resup- (1996) argue that the change in character observed plied with magma (e.g., Dzurisin et al., 1984; Dvorak in the Pu’u ‘O’o eruption occurred because of a and Dzurisin, 1993). Furthermore, if the magma progressive decline in the gas accumulation rate in chamber were isolated in this way through the course E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 97 of the Pu’u ‘O’o–Kupaianaha eruption this would be excellent test case because the range of data collect- reflected in the temperature and chemistry of the ed during the eruption is exceptional and the quality erupting lava. Instead, the most recent study of the of the data is extremely good. The eruption provides, Pu’u ‘O’o–Kupaianaha eruption (Thornber, 2003) therefore, a unique opportunity to examine a Hawai- reinforces the idea that resupply of magma from the ian eruption sequence in great detail. Over the past mantle has occurred throughout the eruption and f10–15 years, the CF model has come to be the shows that what long-term changes have occurred in more widely accepted model of the dynamics of the eruption temperature and magma composition Hawaiian eruptions (e.g., Sparks et al., 1994; Verg- represent the increasing involvement of mantle mag- niolle and Mangan, 2000). Parfitt and Wilson (1994) ma—the exact opposite trend to that which would be pointed out general problems with the model and I seen if the magma chamber were isolated. Thus, all have detailed in this paper the ways in which the CF the available evidence contradicts the idea of an model is inconsistent with the observations made isolated magma chamber and a decline in the supply during the Pu’u ‘O’o eruption, the eruption which of gas to the eruption. the authors of the CF model elected to use as their Parfitt and Wilson (1994) argued that the change in test case (Vergniolle and Jaupart, 1990; Vergniolle, character from intermittent fountaining to continuous 1996). Furthermore, I have shown that the RSD eruption represents a long-term evolution of the dike model, when applied to the same eruption, produces geometry and thermal state, consistent with the obser- results that are consistent with a wide range of vations and interpretation of Dvorak and Okamura observations. My point is not that the CF model is (1985). The change from high fountaining to minor inherently flawed but, instead, that any model has explosive activity, gas release and lava outpouring is value only if it actually reproduces the key features seen as being a result of the decrease in magma of the system under examination. In the case of the eruption rate and rise speed that accompanied the CF model as applied to Hawaiian eruptions, the change from intermittent to continuous eruption. That model is inconsistent in many ways with the obser- the eruption rate was lower during the continuous vational evidence. phase is indisputable. Observations made during the high fountaining phases at Pu’u ‘O’o and during lava outpouring at Kupaianaha show that the typical vol- 7. Transitional eruptions ume flux during high fountaining was f7.7 106 m3 1 6 3  1 dayÀ compared with f0.35 10 m dayÀ at Some basaltic eruptions exhibit behaviour that  Kupaianaha (Heliker et al., 2003). Parfitt and Wilson appears to display features of both Hawaiian and (1994) used these fluxes to estimate the magma rise Strombolian activity and are referred to as ‘‘transi- 1 speed as being f0.3 m sÀ during the high fountain- tional’’ eruptions (Parfitt and Wilson, 1995). The 1973 1 ing episodes and f0.01 m sÀ during the Kupaianaha eruption of Heimaey in Iceland is an example of this eruption. Parfitt and Wilson (1995) have shown that, type of event (see above). Another example is the 6th for magma gas contents and viscosities observed to 29th July 1975 stage of the Great Tolbachik Fissure during the Pu’u ‘O’o–Kupaianaha eruption, the eruption. This eruption is described as Strombolian– RSD model predicts that at a rise speed of 0.3 m Plinian by Maleyev and Vande-Kirkov (1983). They 1 sÀ , the activity should be Hawaiian, and at a rise say that the eruption ‘‘ejected a continuous stream 1 speed of 0.01 m sÀ , the activity should be Strombo- of pyroclastic material to a height of 8–11 km’’. lian, consistent with the observed change in eruption Tokarev (1983) describes the eruption as a ‘‘non-stop character. vertical jet of incandescent gases, ash, cinder and volcanic bombs’’ that reached ‘‘a height of 1–1.5 6.5. Discussion km, while above it, to a height of 6–8 km, rose a billowing cloud of ash blown sideways by the wind’’. I have discussed this one eruption in detail for Although there were pulsations in the eruption jet, several reasons. Both the RSD and CF models have eruption was continuous (Tokarev, 1983). Clasts up to been tested using the Pu’u ‘O’o eruption. It is an 2–3 m in diameter were produced and accumulation 98 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 of this material around the vent generated a cinder because the short time gap between Strombolian cone but did not produce any lava flows (Maleyev and explosions means that from the point of view of the Vande-Kirkov, 1983; Tokarev, 1983). Thus, this erup- heat output the activity is continuous and can thus tion, like the Heimaey eruption, exhibits character- generate a sustained plume. The height of the plume is istics of both Hawaiian and Strombolian activity. much greater than that associated with Hawaiian In addition to exhibiting characteristics of both eruptions and this difference is expected to be related Hawaiian and Strombolian styles, basaltic eruptions to the difference in grainsize of the erupted material frequently exhibit rapid transitions between these end- compared with a pure Hawaiian eruption (Parfitt and member types of activity. For example, Bertagnini et Wilson, 1999). al. (1990) described such behaviour during the 1989 The CF model has not been used to look at eruption of Etna. They describe how each ‘‘eruptive transitional eruptions or to explain how changes in episode began with a weak strombolian activity, with gas accumulation rates or magma viscosity can lava clasts thrown just beyond the crater rim.’’ As the account for the types of rapid transition in eruption magma level rose in the vent, the explosions became style which are a common feature of basaltic activity. ‘‘more frequent and more violent’’ until they were As we have seen, the sudden change in eruption ‘‘nearly continuous’’. This activity then evolved into character which occurred during the Pu’u ‘O’o– activity which was ‘‘typically hawaiian, with lava Kupaianaha eruption is explained in the CF model fountains up to 100–200 m in height’’ and which by a gradual change in gas accumulation which generated lava flows. occurs over a period of years. Changes in character Parfitt and Wilson (1995) used the RSD model to from Strombolian to Hawaiian and back again, like investigate the nature of transitional eruptions and the those described at Etna, can occur on time scales of conditions which give rise to them. The results of this only hours. modelling (Fig. 5) show that transitional activity is expected to arise primarily when the magma rise speed is intermediate between that of Hawaiian and 8. Conclusions Strombolian eruptions (Table 1) and furthermore that gradual changes in rise speed will give rise to a During the past 20 years, two very different progressive change in eruption character from Strom- models have been proposed to explain the dynamics bolian to Hawaiian or vice versa. The time frame over of explosive basaltic eruptions—the rise speed de- which the eruption character changes is then a func- pendent model (Wilson, 1980; Wilson and Head, tion of the rate at which the magma rise speed 1981; Head and Wilson, 1987; Fagents and Wilson, changes. The modelling suggests that, for example, 1993; Parfitt and Wilson, 1994, 1999; Parfitt et al., as magma rise speed increases the eruption character 1995) and the collapsing foam model (Vergniolle and would change from widely spaced Strombolian explo- Jaupart, 1986, 1990; Jaupart and Vergniolle, 1988, sions to more frequent explosions with the strength of 1989; Vergniolle, 1996). Both models are in current the explosions being fairly constant. Then as the rise usage, often without acknowledgement that an alter- speed increases further the explosions will become native model exists. In this paper, I have examined more closely spaced in time still and will rapidly the basic assumptions made in each model and increase in violence throwing clasts much higher in shown that neither model is flawed in any funda- the air. Continued increase in rise speed then gives mental way, i.e., that each model could apply in a rise to continuous high lava fountaining activity. This given set of conditions. The purpose of a model is, pattern of behaviour is remarkably similar to that however, to represent some behaviour that we ob- described above for the 1989 Etna eruption. serve in nature. Thus, the value of any model The model developed by Parfitt and Wilson (1995) depends on how well it can reproduce this real does not explicitly look at the behaviour of the finer behaviour. Volcanologists examining explosive ba- material ejected in the eruption. A characteristic of saltic activity are at a considerable advantage com- many transitional eruptions is the high sustained pared with those interested in more violent, silicic, eruption plume they develop. Presumably, this arises events in that basaltic explosions occur frequently E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 99 and are relatively safe to monitor. Thus, large bodies exist beneath the vent in order to operate whereas the of literature and data are available to test the two processes invoked by the RSD model can occur in any models. I have examined the way in which each open system as long as the rise speed of the magma is model has been used to explain Strombolian, Hawaii slow or negligible. and transitional eruption styles, and in doing so have arrived at the following conclusions. 8.2. Hawaiian eruptions

8.1. Strombolian eruptions In the case of Hawaiian activity the RSD and CF models present very different and mutually exclusive Both models agree that Strombolian eruptions views of the eruption dynamics. The models make result from the accumulation and bursting of a gas fundamentally different assumptions about the flow pocket at shallow depths within an open magmatic regime which prevails during eruption: The RSD system. This is consistent with direct observations model assumes homogeneous two-phase flow and (e.g., Vergniolle and Brandeis, 1994) and with acous- the CF model assumes segregated two-phase flow. tic wave studies (Vergniolle and Brandeis, 1994, They also assume different ‘driving’ gases—in typical 1996; Vergniolle et al., 1996). The models diverge, conditions, the RSD model will only work if H2O is however, in the assumptions they make about where the dominant gas and the CF model only works if CO2 the gas segregates from the magma. In the RSD is the driving gas. My examination of the two models model segregation is thought to be progressive and suggests that neither model is fundamentally flawed occurs because of the low rise speed of the magma and thus that either might operate in different starting beneath the eruptive vent. Such a model is consistent conditions. with observational evidence (Table 1) which shows Both models have been tested on the same eruption that Strombolian eruptions are associated with low (the 1983–1988 Pu’u ‘O’o eruption of Kilauea vol- magma rise speeds. The CF model assumes that gas cano). I have examined in detail how well each model segregation occurs at depth in a magma chamber or fits with the behaviour observed during this eruption storage zone and that accumulation of this gas as a and have shown that the CF model (Vergniolle and foam layer and its partial collapse give rise to a slug Jaupart, 1990; Vergniolle, 1996) fails to explain many of gas which rises up through the vent system and key aspects of the eruption. For instance, observations bursts through the top of the magma column. Thus, show that the dominant gas released in the eruption is the CF model requires a special set of conditions to H2O (85%) and only 3% of the released gas is CO2 exist whereas the RSD model is applicable to any which the CF model assumes is the driving gas. open system. Furthermore, observational evidence does not support Recent studies of seismicity at Stromboli (Neuberg the contention that gas segregation and concentration et al., 1994; Chouet et al., 1999) show that earth- has occurred at depth prior to eruption but is instead quakes are generated in direct association with each consistent with the homogeneous flow assumed in the Strombolian explosion. The source of such earth- RSD model. The RSD model can also explain many quakes is located several hundred metres beneath other aspects of the eruptive behaviour such as the the surface. It has been suggested that these earth- characteristic deflation pattern observed during each quakes are caused by the collapse and movement of eruption, the long-term changes in the duration and the gas slug at depth. Further modelling work is eruption rates observed during the eruption sequence, needed to show that the seismic waves generated in and the change in eruption character which occurred Strombolian explosions are generated in this way but in 1986. such work provides a potential way to determine The problems highlighted by the application of whether gas accumulation and foam formation occurs the CF model to the Pu’u ‘O’o eruption are at Stromboli. It should be stressed, though, that there sufficiently far-reaching that they draw the validity is no reason why both models should not be applica- of the model in its application to other Hawaiian ble in different systems and it must be understood that eruptions into serious question. I would therefore the CF model requires a particular set of conditions to urge considerable caution in the use of this model 100 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 in treating the dynamics of Hawaiian eruptions. It character can occur rapidly in response to changes must also be borne in mind, though, that the models in rise speed. By contrast, in the CF model, eruption have only been tested in detail on this one eruption. character is determined by magma viscosity and gas While the evidence from the Pu’u ‘O’o eruption accumulation rates (Jaupart and Vergniolle, 1988; supports the RSD model as the most valid treat- Vergniolle, 1996), and it is hard to see how the rapid ment, more testing on other Hawaiian eruptions is transitions observed in basaltic eruptions can be needed to resolve the issue of which model is most explained by such a model. applicable to Hawaiian eruptions and, indeed, to explosive basaltic eruptions in general. It should be noted, though, that like with Strombolian eruptions Acknowledgements the CF model requires a particular set of conditions to prevail beneath the vent in order to be applicable I thank Christina Heliker for comments on the (a large enough storage zone at a depth where gas origin of data used in Vergniolle and Jaupart (1990).I is exsolving) whereas the RSD model has no such also thank the staff of the Hawaiian Volcano limitations. Observatory, especially Tom Wright, for providing access to data used here (and elsewhere) and for 8.3. Transitional eruptions discussion of many of the issues raised in this paper. Thanks to Andy Harris for comments relating to I have noted that some explosive basaltic eruptions activity at Etna. This paper has benefited from show features of both Hawaiian and Strombolian detailed reviews by Sylvie Vergniolle, Greg Valentine eruption styles and these have been denoted ‘‘transi- and Roberto Scandone. Finally, thanks to Lionel tional’’ eruptions by Parfitt and Wilson (1995). In Wilson for many, many discussions of these issues addition, basaltic eruptions frequently show rapid over the years. transitions in character from Strombolian to Hawaiian and vice versa. The RSD model was used by Parfitt and Wilson (1995) to investigate the conditions in Appendix A. Treatment of simultaneous magma which transitional eruption styles arise. They sug- eruption and drainback in conduit flow gested that transitional eruptions arise when the magma rise speed is too high to yield purely Strom- A.1. Introduction bolian activity and too low to yield purely Hawaiian behaviour. For a typical basaltic eruption this transi- It is assumed that all of the magma behaves as a tion occurs in the magma rise speed range 0.01–0.1 Newtonian fluid with the same viscosity and that for 1 msÀ (Fig. 5). This is consistent with a range of both upward and downward magma streams there is observational data (e.g., Table 1). The model further no variation with depth of the magma density and the suggests that as magma rise speed progressively pressure gradient driving the motion. There is no increases or decreases, an eruption can rapidly change guarantee that the pressure gradients in real volcanic in character from Strombolian to Hawaiian or vice conduits are independent of depth. However, as versa. In the case of increasing rise speed, for exam- shown by Wilson et al. (1980), Wilson and Head ple, it would be expected that Strombolian explosions (1981) and Giberti and Wilson (1990), for mafic will become progressively more frequent but with magmas, there exists a wide range of possible erup- little change in violence until a rapid change occurs tion conditions in which this condition is approxi- during which explosions become very closely spaced mately satisfied, even when rising magma exsolves in time and much more violent. This behaviour then sufficient volatiles to undergo fragmentation. The gives way to continuous lava fountaining. This kind main restriction on the application of the following of transition in character is very similar to that calculations, therefore, is the assumption of constant observed in many real basaltic eruptions. In the magma densities. There are two cases to be consid- RSD model, the primary control on eruption character ered. In the first, both the rising and the descending is the magma rise speed and thus transitions in magma are unvesiculated and their motions are E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 101 laminar. In the second, the rising magma is assumed To find the average velocity uav-up of the upward- to have vesiculated to the point of fragmentation into moving fluid, we weight the local value of u by the pyroclasts and released gas, and the upward motion relative volume of fluid with that velocity, (2prdr), is turbulent. so that:

r r A.2. Laminar upward motion 0 0 uav up 2purdr 2prdr A4 À ¼ Z0 Z0 ð Þ This analysis can be applied directly to the which yields: motion of volatile-poor magma in a conduit where there is either net drainback from a lava pond around 2 the vent or net growth of such a pond. It can also be uav up dP=dx up= 8l r0: A5 À ¼ ½ð Þ ð ފ ð Þ inferred to apply to conditions in which vesiculating magma is erupting into a lava fountain while The upward mass flux Mup is then: degassed magma is draining back into the conduit 2 Mup p r0uav upqup; A6 system provided that the values derived are taken to ¼ À ð Þ refer to conditions at depths greater than a few where qup is the bulk density of the upward-flowing hundred metres, where variations in bulk magma fluid. density are still small, there is negligible upward Next consider the fluid moving downward. Skel- acceleration of the rising magma, and similar viscos- land (1967, Chap. 3) derives equations for flow in an ities can be assumed for both the rising and the annulus of a Bingham plastic fluid, and by setting the sinking magma. yield strength of such a fluid equal to zero the velocity The basic relationship controlling flow of a New- profile u(r) is found to be: tonian fluid in a circular conduit is: u dP dx 4 r2 r2 r2 r2 l du=dr 1=2 r dP=dx ; A1 = down= l w w 0 ¼ Àð Þ ð Þ ð Þ ¼ ½ð Þ ð ފf À þ ½ð À Þ =ln rw=r0 ln r=rw : A7 where l is the fluid viscosity, (dP/dx) is the pressure ð ފ ð Þg ð Þ gradient driving the motion in the x direction, r is the 2 2 Note that u is zero at r =rw because then (rw r )=0 radial co-ordinate and u is the local velocity of the À and ln(r/rw)=ln(1)=0; also u is zero at r =r0 because fluid. We assume that between r =0 and r =r , where r 0 0 then ln(r/rw)=ln(r0/rw)= ln(rw/r0). The maximum ve- is some intermediate radius, the motion is upward and À locity ua must occur at the radius ra for which du/ between r =r0 and r =rw, where rw is the radius of the dr =0. Differentiating Eq. (A7): confining wall, the motion is downward. Thus, u =0 at both r =r0 and at r =rw. du=dr dP=dx = 4l 2r r2 r2 First, consider the fluid moving upward. Integrat- ¼ ½ð Þdown ð ފfÀ þ ½ð w À 0Þ ing Eq. (A1) between r =0 and a general value of r =ln rw=r0 1=r A8 ð ފð Þgð Þ gives the velocity profile u(r): and setting the term in curly brackets in Eq. (A8) to zero gives: u dP=dx = 4l r2 r2 : A2 ¼ ½ð Þup ð ފ ð 0 À Þ ð Þ 2 2 2 2 2 2 ra rw r0 = 2 ln rw=r0 rw r0 =ln rw=r0 : In this equation and those that follow, (dP/dx)up ¼ð À Þ ½ ð ފ ¼ ð À Þ ð Þ A9 represents the absolute value of the pressure gradient; ð Þ the driving pressure must of course decrease in the direction of motion for the velocity to be positive. The The value of the maximum velocity, ua, is given by: maximum velocity uc occurs at r =0 and is: 2 2 2 2 ua dP=dx down= 4l rw ra rw r0 2 ¼ ½ð Þ ð ފf À þ ½ð À Þ uc dP=dx = 4l r : A3 =ln rw=r0 ln ra=rw : A10 ¼ ½ð Þup ð ފ 0 ð Þ ð ފ ð Þg ð Þ 102 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107

Table A1 and the average velocity of the fluid, uav-down, is obtained using the equivalent of Eq. (A4) as: Values of the ratios of upward to downward pressure gradient, magma mass flux, maximum fluid velocity and mean fluid velocity 2 2 2 2 as a function of the fraction of the conduit radius occupied by the uav down dP=dx down= 8l rw r0 rw r0 upward flow for the case where the upward motion is laminar and À ¼ ½ð Þ ð ފf þ À ½ð À Þ [ / ]=0.8 =ln rw=r0 : A11 qup qdown ð ފg ð Þ r0/rw (dP/dx)up / Mup / uc /ua uav-up / (dP/dx)down Mdown uav-down The downward mass flux Mdown is then: 0.154 10.00 0.092 0.611 0.473 2 2 0.232 5.00 0.152 0.870 0.666 Mdown p rw r0 uav downqdown; A12 ¼ ð À Þ À ð Þ 0.386 2.00 0.327 1.543 1.169 where qdown is the bulk density of the downward- 0.534 1.00 0.621 2.589 1.950 flowing fluid. 0.642 0.61 1.000 3.892 2.925 The final issue is to establish how the boundary 0.683 0.50 1.214 4.621 3.471 0.838 0.20 3.007 10.646 7.988 between the two flows at r =r0 is related to the other 0.910 0.10 6.004 20.656 15.494 variables. This involves looking at the continuity of the overall velocity profile. We have already ensured that the velocity itself is continuous by arranging that the amounts by which the total pressure gradients u=0 at r =r0 in both Eq. (A2) for the upward velocity differs from the static weights of the magma in each and Eq. (A7) for the downward velocity. We now case. Assume that the pressure in the magma reser- require that the slopes of the two functions also be voir at depth is the hydrostatic weight of the over- continuous. For the upward velocity profile, Eq. (A2) lying crust (in practice, the reservoir is likely to be leads to: overpressured relative to the lithostatic load by a few MPa, but this does not significantly affect the du=dr dP=dx = 4l 2r A13 ¼ ½ð Þup ð ފðÀ ÞðÞ following illustration) and that the upward magma and we have already obtained the derivative of the flow occurs because the magma in the centre of the downward velocity profile as Eq. (A8). Equating the conduit is positively buoyant. Similarly, the magma in the outer, descending annulus is assumed to be two at r =r0 and simplifying: negatively buoyant. If the crustal density is qcrust, we dP=dx = dP=dx r2 =r2 1 then have: ð Þup ð Þdown ¼f½ð w 0ÞÀ Š 2 =ln rw=r0 1; A14 ð Þ gÀ ð Þ dP=dx g q q A16 ð Þup ¼ ð crust À upÞðÞ so for any choice of the ratio (rw/r0), Eq. (A14) gives the ratio of [(dP/dx) /(dP/dx) ] that ensures con- up down dP=dx g q q A17 tinuity of the velocity profile. Then taking the ratio of ð Þdown ¼ ð down À crustÞðÞ Eqs. (A6) and (A12): 3 Plausible values might be qcrust=2250 kg mÀ , qup= 3 3 2000 kg mÀ and =2500 kg mÀ , in which case Mup=Mdown qup=qdown dP=dx up= dP=dx down qdown ¼½ Š½ð Þ ð Þ Š [q /q ]=0.8, as illustrated in Table A1. The table r4= r2 r2 r2 r2 r2 r2 up down ½ 0 fð w À 0Þf w þ 0 þ ½ð w À 0Þ shows that for the case of pure convection, where =ln rw=r0 : A15 M /M =1, r /r =0.642, which implies that ð ފggŠ ð Þ up down 0 w f41% of the area of the conduit is occupied by Thus, given a choice of the density ratio [qup/ rising magma and f59% by descending magma. qdown], the ratio of the mass fluxes can be obtained. Clearly, it is necessary that qup is always less than Table A1 shows a selection of examples, tabulated as qdown, and depending on the amount of gas exsolv- a function of (r0/rw) for [qup/qdown]=0.8. The reason ing from the rising magma in the deep part of the for this choice of [qup/qdown] is as follows. The conduit, and on the crustal density, likely ranges of pressure gradients driving the upward and downward values are f0.55 to 0.85 for [qup/qdown] and f0.3 movement of the magmatic fluids are by definition to at least 3 for [(dP/dx)up/(dP/dx)down]. E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 103

A.3. Turbulent upward motion Equating this to Eq. (A21) with r=r0, taking account of the fact that the upward and downward This treatment applies to the shallow part of a velocities have opposite signs: conduit where enough volatile exsolution has oc- curred that magma fragmentation had taken place. 2 uav up=uav down 2 ln rw=r0 rw=r0 1 Both the viscosity and the density of the magma are À À ¼f½À ð Þþð Þ À Š 1 a=2 b=17 = a 16b very different in the upward and downward magma ½ À À Šg f½ þ Š 2 streams. The velocity profile in the outer annulus is rw=r0 1 ln rw=r0 Âf½ð Þ þ Š ð Þ the same as in the previous section but in the central 2 rw=r0 1 : A23 region out to radius r0 the velocity profile is of the Àð Þ þ gg ð Þ form (Knudson and Katz, 1958): Then taking the ratio of Eqs. (A6–12) and substi- 2 32 tuting Eq. (A23) for (uav-up/uav-down): u uc 1 a r=r0 b r=r0 A18 ¼ ½ À ð Þ À ð Þ ŠðÞ where a=0.204 and b=0.796. The average velocity Mup=Mdown q =q 1 a=2 b=17 u is found from Eq. (A4): ¼ð up downÞ½ð À À Þ av-up 2 = a 16b 2 ln rw=r0 rw=r0 ð þ ފ½À ð Þþð Þ r r 2 0 0 1 = rw=r0 1 ln rw=r0 uav up 2purdr 2prdr À Š ff½ð Þ þ Š ð Þ À 2 2 ¼ Z0 Z0 rw=r0 1 rw=r0 1 : A24 2 2 4 2 Àð Þ þ g½ð Þ À Šg ð Þ 2uc =r r =2 ar = 4r ¼ ½ð Þ 0Š À ð 0Þ Â r0 34 32 Thus, for any choice of the ratio (rw/r0), the ratio br 34r0 À  Z0  of the upward and downward average velocities can be found from Eq. (A23) and of the upward and and so: downward mass fluxes can be obtained from Eq. (A24). Some values of (uav-up/uav-down) are given in uav up uc 1 a=2 b=17 0:8512 uc: A19 À ¼ ½ À À Š¼ ð Þ Table A2. However, in order to specify values of The first derivative of the velocity profile is:

Table A2 2 31 32 du=dr uc 2ar=r0 32br =r0 A20 Values of the ratios of upward to downward magma mass flux and ¼ ½À À ŠðÞ mean magma velocity as a function of the fraction of the conduit and using Eq. (A19) to write this in terms of u , radius occupied by the upward flow for the case where the upward av-up motion is turbulent and [ / ]=0.024 becomes: qup qdown r0 /rw Mup /Mdown uav-up /uav-down 2 31 32 0.10 0.000011 0.046 du=dr uav up 2ar=r0 32br =r0 ¼ À ½À À Š 0.20 0.000077 0.076 = 1 a=2 b=17 : A21 0.30 0.000269 0.113 ½ À À Š ð Þ 0.40 0.000738 0.161 0.50 0.001823 0.228 This must now be equated to the first derivative of 0.60 0.004417 0.327 the annulus flow, Eq. (8), at r=r0. It is convenient to 0.70 0.011349 0.492 use Eq. (A11) to eliminate [(dP/dx)down/(4l)] from 0.80 0.035050 0.821 Eq. (A8) giving: 0.90 0.185045 1.808 0.95 0.840210 3.782 0.96 1.345392 4.769 2 2 2 2 0.97 2.450476 9.702 du=dr 2uav down= rw r0 rw r0 ¼½ À f þ À ½ð À Þ 0.99 23.13130 19.569 2 2 =ln rw=r0 2r rw r0 0.999 2362.51119 197.171 ð ފgŠÂfÀ þ ½ð À Þ 0.9999 236761.95087 1973.312 =ln rw=r0 1=r : A22 ð ފð Þg ð Þ 104 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107

Aramaki, S., Hayakawa, Y., Fujii, T., Nakamura, K., Fukuoka, T., (Mup/Mdown), it is necessary to decide on a plausible value of (q /q ). The fragmented magma in the 1986. The October 1983 eruption of Miyakejima volcano. up down J. Volcanol. Geotherm. Res. 29, 203–229. rising part of the conduit will have a bulk density Bertagnini, A., Calvari, S., Coltelli, M., Landi, P., Pompilio, M., qup given by: Scribano, V., 1990. The 1989 eruptive sequence. In: Barberi, F., Bertagnini, A., Landi, P. (Eds.), Mt. Etna: the 1989 eruption, C.N.R.–Gruppo Nazionale Per La Vulcanologia Italy, Giardini. 1=qup n=qg 1 n =qmelt A25 ð Þ¼ð Þ þ ½ð À Þ ŠðÞ Blackburn, E.A., Wilson, L., Sparks, R.S.J., 1976. Mechanisms and where n is the mass fraction of exsolved volatile, q dynamics of Strombolian activity. J. Geol. Soc. (Lond.) 132, g 429–440. is the gas density and qmelt is the density of unvesi- Carracedo, J.C., Rodriguez Badiola, E., Soler, V., 1992. The 1730– culated rising magma. Consider a typical basaltic 1736 eruption of Lanzarote, Canary Islands: a long, high mag- melt having exsolved 0.25 wt.% H2O, i.e., a mass nitude basaltic fissure eruption. J. Volcanol. Geotherm. Res. 53, fraction n=0.0025; at a typical magmatic temperature 239–250. of 1450 K, the density of this gas emerging from the Cas, R.A.F., Wright, J.V., 1988. Volcanic Successions. Chapman & Hall, London, p. 528. vent, where the pressure is close to atmospheric, is Cashman, K.V., Mangan, M.T., 1995. Physical aspects of magmatic 3 3 f0.15 kg mÀ . Thus, qup is f60 kg mÀ . Using degassing: II. Constraints on vesiculation processes from textur- 3 qdown=2500 kg mÀ , as before, (qup/qdown)=0.024, al studies of eruptive products. In: Carroll, M. (Ed.), Volatiles in Magma. Mineral. Soc. Am., Washington, DC, pp. 447–478. and this value is used to generate the values of (Mup/ M ) in Table A2. Cattermole, P., 1989. Planetary Volcanism: A Study of Volcanic Ac- down tivity in the Solar System. Ellis Horwood, Chichester, pp. 1–443. It is much less easy in this case to specify the ratios Chouet, B., Hamisevicz, N., McGetchin, T.R., 1974. Photoballistics of the maximum velocities and the pressure gradients. of volcanic jet activity at Stromboli, Italy. J. Volcanol. Geo- The normal treatment for turbulent flow is to relate the therm. Res. 32, 4961–4976. pressure gradient to the mean velocity by: Chouet, B., Saccorotti, G., Dawson, P., Martini, M., Scarpa, R., De Luca, G., Milana, G., Cattaneo, M., 1999. Broadband measure- 2 ments of the sources of explosions at Stromboli volcano, Italy. r0 dP=dx up f qupuav up A26 ð Þ ¼ À ð Þ Geophys. Res. Lett. 26, 1937–1940. 3 where f is a dimensionless friction factor of order 10À Crisp, J.A., 1984. Rates of magma emplacement and volcanic out- for the case of motion of a fluid in a pipe with smooth put. J. Volcanol. Geotherm. Res. 20, 177–211. Dvorak, J.J., Dzurisin, D., 1993. Variations in magma supply rate at walls (probably the best assumption in this case). Kilauea volcano, Hawaii. J. Geophys. Res. 98, 22255–22268. However, Eq. (A26) assumes that there is no acceler- Dvorak, J.J., Okamura, A.T., 1985. Variations in tilt rate and har- ation of the magma, whereas if the rising magma is monic tremor amplitude during the January–August 1983 East fragmenting to form a lava fountain much of the Rift eruptions of Kilauea volcano, Hawaii. J. Volcanol. Geo- ambient pressure decrease is used to accelerate the therm. Res. 25, 249–258. Dzurisin, D., Koyanagi, R.Y., English, T.T., 1984. Magma supply magma and very little is used to overcome wall and storage at Kilauea volcano, Hawaii, 1956–1983. J. Volca- friction. Thus, the value needed for (dP/dx)up in Eq. nol. Geotherm. Res. 21, 177–206. (A26) would be very much less than the value used in Eaton, J.P., Richter, D.H., Krivoy, H.L., 1987. Cycling magma the slow-speed, laminar flow case and, furthermore, between the summit reservoir and Kilauea Iki lava lake during would change with depth below the surface. No the 1959 eruption of Kilauea volcano. U. S. Geol. Surv. Prof. Pap. 1350, 1307–1335. attempt is made to take the analysis further here, but Fagents, S.A., Wilson, L., 1993. Explosive volcanic eruptions: VII. Table A2 shows that a given ratio of Mup/Mdown The range of pyroclasts ejected in transient explosions. Geo- occurs much closer to the conduit wall in the case phys. J. Int. 113, 359–370. of turbulent upward flow than in the laminar case. Fedotov, S.A., Kovalev, G.N., Markhinin, Y.K., Slezin, Y.B., Tsyur- upa, A.I., Gusev, N.A., Andreyev, V.I., Leonov, V.L., Ovsyan- nikov, A.A., 1983. Chronology and features of the Southern Breakthrough of the Great Tolbachik fissure eruption 1975– References 1976. In: Fedotov, S.A., Markhinin, Y.K. (Eds.), The Great Tolbachik Fissure Eruption: Geological and Geophysical Data

Andres, R.J., Kyle, P.R., Stokes, J.B., Rose, W.I., 1989. SO2 from 1975–76. Cambridge Univ. Press, Cambridge, pp. 11–26. episode 48A eruption, Hawaii: sulfur dioxide emissions from Gerlach, T.M., 1986. Exsolution of H2O, CO2, and S during erup- the episode 48A East Rift Zone eruption of Kilauea volcano, tive episodes at Kilauea volcano, Hawaii. J. Geophys. Res. 91, Hawaii. Bull. Volcanol. 52, 113–117. 12177–12185. E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 105

Gerlach, T.M., Graeber, E.J., 1985. Volatile budget of Kilauea vol- Jaupart, C., Vergniolle, S., 1988. Laboratory models of Hawaiian cano. Nature 313, 273–277. and Strombolian eruptions. Nature 331, 58–60. Gerlach, T.M., Taylor, B.E., 1990. Carbon isotope constraints on Jaupart, C., Vergniolle, S., 1989. The generation and collapse of degassing of carbon dioxide from Kilauea volcano. Geochim. foam layer at the roof of a basaltic magma chamber. J. Fluid Cosmochim. Acta 54, 2051–2058. Mech. 203, 347–380. Giberti, G., Wilson, L., 1990. The influnce of geometry on the Klein, F.W., Koyanagi, R.Y., Nakata, J.S., Tanigawa, W.R., 1987. ascent of magma in open fissures. Bull. Volcanol. 52, 515–521. The seismicity of Kilauea’s magma system. U. S. Geol. Surv. Giberti, G., Jaupart, C., Sartoris, G., 1992. Steady-state operation of Prof. Pap. 1350, 1019–1185. Stromboli volcano, Italy: constraints on the feeding system. Knudson, J.G., Katz, D.L., 1958. Fluid Dynamics and Heat Trans- Bull. Volcanol. 54, 535–541. fer. McGraw-Hill, New York. 576 pp. Greenland, L.P., 1984. Gas composition of the January 1983 erup- Lockwood, J.P., Dvorak, J.J., English, T.T., Koyanagi, R.Y., Oka- tion of Kilauea volcano, Hawaii. Geochim. Cosmochim. Acta mura, A.T., Summers, M.L., Tanigawa, W.R., 1987. Mauna Loa 48, 193–195. 1974–1984: a decade of intrusive and extrusive activity. U. S. Greenland, L.P., 1988. Gases from the 1983–84 east rift eruption. Geol. Surv. Prof. Pap. 1350, 537–570. U. S. Geol. Surv. Prof. Pap. 1463, 145–153. Maleyev, Y.F., Vande-Kirkov, Y.V., 1983. Features of pyroclastics Greenland, L.P., Rose, W.I., Stokes, J.B., 1985. An estimate of gas of the Northern Breakthrough of the Great Tolbachik Fissure emissions and magmatic gas content from Kilauea volcano. Eruption and the origin of its pale-grey ash. In: Fedotov, S.A., Geochim. Cosmochim. Acta 49, 125–129. Markhinin, Y.K. (Eds.), The Great Tolbachik Fissure Eruption. Harris, A.J.L., Neri, M., 2002. Volumetric observations during par- Cambridge Univ. Press, Cambridge, pp. 57–71. oxysmal eruptions at Mount Etna: pressurized drainage of a Manga, M., Stone, H.A., 1994. Interactions between bubbles in shallow chamber or pulsed supply? J. Volcanol. Geotherm. magmas and lavas: effects of bubble deformation. J. Volcanol. Res. 116, 79–95. Geotherm. Res. 63, 267–279. Head, J.W., Wilson, L., 1987. Lava fountain heights at Pu’ u ‘O’o, Mangan, M.T., Cashman, K.V., 1996. The structure of basaltic Kilauea, Hawaii: indicators of amount and variations of scoria and reticulite and inferences for vesiculation, foam for- exsolved magma volatiles. J. Geophys. Res. 92, 13715–13719. mation, and fragmentation in lava fountains. J. Volcanol. Geo- Head, J.W., Wilson, L., 1989. Basaltic pyroclastic eruptions: influ- therm. Res. 73, 1–18. ence of gas-release patterns and volume fluxes on fountain Mangan, M.T., Cashman, K.V., Newman, S., 1993. Vesiculation of structure, and the formation of cinder cones, spatter cones, root- basaltic magma during eruption. Geology 21, 157–160. less flows, lava ponds and lava flows. J. Volcanol. Geotherm. Mangan, M.T., Heliker, C.C., Mattox, T.N., Kauahikaua, J.P., Helz, Res. 37, 261–271. R.T., 1995. Episode 49 of the Pu’u ‘O’o–Kupaianaha eruption Head, J.W., Wilson, L., 2003. Deep submarine pyroclastic erup- of Kilauea volcano—breakdown of a steady-state eruptive era. tions: theory and predicted landforms and deposits. J. Volcanol. Bull. Volcanol. 57, 127–135. Geotherm. Res. 121, 155–193. Mattox, T.N., Heliker, C., Kauahikaua, J., Hon, K., 1993. Devel- Head, J.W., Crumpler, L.S., Aubele, J.C., Guest, J.E., Saunders, opment of the 1990 Kalapana flow field, Kilauea volcano, R.S., 1992. Venus volcanism: classification of volcanic fea- Hawaii. Bull. Volcanol. 55, 407–413. tures and structures, associations, and global distribution from McGetchin, T.R., Ulrich, W.G., 1973. Xenoliths in maars and diat- Magellan data. J. Geophys. Res. 97, 13153–13197. remes with inferences for the Moon, Mars and Venus. J. Geo- Heiken, G., 1972. Morphology and petrology of volcanic ashes. phys. Res. 78, 1833–1853. Geol. Soc. Amer. bull. 83, 1961–1988. McGetchin, T.R., Settle, M., Chouet, B.A., 1974. Cinder cone Heiken, G., 1978. Characteristics of tephra from Cinder Cone, Las- growth modeled after Northeast Crater, Mount Etna, Sicily. sen Volcanic National Park, California. Bull. Volcanol. 41–2, J. Geophys. Res. 79, 3257–3272. 119–130. Neuberg, J., Luckett, R., Ripepe, M., Braun, T., 1994. Highlights Heliker, C., Mattox, T.N., 2003. The first two decades of the Pu’u from a seismic broadband array on Stromboli volcano. Geophys. ‘O’o–Kupaianaha eruption: chronology and selected bibliogra- Res. Lett. 21, 749–752. phy. U. S. Geol. Surv. Prof. Pap. 1676, 1–28. Okamura, A.T., Dvorak, J.J., Koyanagi, R.Y., Tanigawa, W.R., Heliker, C., Wright, T.L., 1991. The Pu’u ‘O’o–Kupaianaha erup- 1988. Surface deformation during dike propagation: the 1983 tion of Kilauea. Eos 73, 521–530. east rift eruption of Kilauea volcano, Hawaii. U. S. Geol. Surv. Heliker, C., Swanson, D.A., Takahashi, T.J., 2003. The Pu’u ‘O’o– Prof. Pap. 1463, 165–181. Kupaianaha eruption of Kilauea Volcano Hawai’i: The first 20 Parfitt, E.A., 1998. A study of clast size distribution, ash deposition years. U.S.G.S. Prof. Pap. 1676, 1–206. and fragmentation in a Hawaiian-style volcanic eruption. J. Vol- Hort, M., Seyfried, R., 1998. Volcanic eruption velocities measured canol. Geotherm. Res. 84, 197–208. with a micro radar. Geophys. Res. Lett. 25, 113–116. Parfitt, E.A., Wilson, L., 1994. The 1983–86 Pu’ u ‘O’o eruption of Houghton, B.F., Schmincke, H.-U., 1989. Rothenberg scoria cone, Kilauea volcano, Hawaii: a study of dike geometry and eruption East Eifel: a complex Strombolian and phreatomagmatic volca- mechanisms for a long-lived eruption. J. Volcanol. Geotherm. no. Bull. Volcanol. 52, 28–48. Res. 59, 179–205. Houghton, B.F., Wilson, C.J.N., 1989. A vesicularity index for Parfitt, E.A., Wilson, L., 1995. Explosive volcanic eruptions: pyroclastic deposits. Bull. Volcanol. 51, 451–462. IX. The transition between Hawaiian-style lava fountaining 106 E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107

and Strombolian explosive activity. Geophys. J. Int. 121, ample of quasi-steady-state magma transfer. U.S.G.S. Prof. Pap. 226–232. 1350, 405–469. Parfitt, E.A., Wilson, L., 1999. A Plinian treatment of fallout Tokarev, P.I., 1983. Calculation of the magma discharge, growth in from Hawaiian lava fountains. J. Volcanol. Geotherm. Res. the height of the cone and dimensions of the feeder channel of 88, 67–75. Crater 1 in the Great Talbachik Fissure Eruption, July 1975. Parfitt, E.A., Wilson, L., Neal, C.A., 1995. Factors influencing the In: Fedotov, S.A., Markhinin, Y.K. (Eds.), The Great Tolbachik height of Hawaiian lava fountains: implications for the use of Fissure Eruption. Cambridge Univ. Press, pp. 27–35. fountain height as an indicator of magma gas content. Bull. Vergniolle, S., 1996. Bubble size distribution in magma chambers Volcanol. 57, 440–450. and dynamics of basaltic eruptions. Earth Planet. Sci. Lett. 140, Proussevitch, A.A., Sahagian, D.L., 1996. Dynamics of coupled 269–279. diffusive and decompressive bubble growth in magmatic sys- Vergniolle, S., Brandeis, G., 1994. Origin of sound generated by tems. J. Geophys. Res. 101, 17447–17455. Strombolian explosions. Geophys. Res. Lett. 21, 1959–1962. Richter, D.H., Eaton, J.P., Murata, K.J., Ault, W.U., Krivoy, H.L., Vergniolle, S., Brandeis, G., 1996. Strombolian explosions: 1. A 1970. Chronological narrative of the 1959–60 eruption of large bubble breaking at the surface of a lava column as a source Kilauea volcano, Hawaii. U. S. Geol. Surv. Prof. Pap. 537-E, of sound. J. Geophys. Res. 101, 20433–20447. 1–73. Vergniolle, S., Jaupart, C., 1986. Separated two-phase flow and Ripepe, M., 1996. Evidence for gas influence on volcanic seismic basaltic eruptions. J. Geophys. Res. 91, 12842–12860. signals recorded at Stromboli. J. Volcanol. Geotherm. Res. 70, Vergniolle, S., Jaupart, C., 1990. Dynamics of degassing at Kilauea 221–233. volcano, Hawaii. J. Geophys. Res. 95, 2793–2809. Self, S., 1976. The recent volcanology of Terceira, Azores. J. Geol. Vergniolle, S., Mangan, M.T., 2000. Hawaiian and Strombolian Soc. (Lond.) 132, 645–666. eruptions. In: Sigurdsson, H., Houghton, B., McNutt, S.R., Self, S., Sparks, R.S.J., Booth, B., Walker, G.P.L., 1974. The 1973 Rymer, H., Stix, J. (Eds.), Encyclopedia of Volcanoes. Academ- Heimaey Strombolian scoria deposit, Iceland. Geol. Mag. 111, ic Press, San Diego, CA, pp. 447–461. 539–548. Vergniolle, S., Brandeis, G., Mareschal, J.-C., 1996. Strombolian Seyfried, R., Freundt, A., 2000. Experiments on conduit flow and explosions: 2. Eruption dynamics determined from acoustic eruption behaviour of basaltic volcanic eruptions. J. Geophys. measurements. J. Geophys. Res. 101, 20449–20466. Res. 105, 23727–23740. Wadge, G., 1981. The variation of magma discharge during basaltic Skelland, A.H.P., 1967. Non-Newtonian Flow and Heat Transfer. eruptions. J. Volcanol. Geotherm. Res. 11, 139–168. Wiley, New York. 469 pp. Walker, G.P.L., 1973. Explosive volcanic eruptions—a new classi- Sparks, R.S.J., 1978. The dynamics of bubble formation and growth fication scheme. Geol. Rundsch. 62, 431–446. in magmas: a review and analysis. J. Volcanol. Geotherm. Res. Walker, G.P.L., Self, S., Wilson, L., 1984. Tarawera 1886, New 3, 1–37. Zealand—a basaltic Plinian fissure eruption. J. Volcanol. Geo- Sparks, R.S.J., Barclay, J., Jaupart, C., Mader, H., 1994. Physical therm. Res. 21, 61–78. aspects of magma degassing: I. Experimental and theoretical Wallace, P.J., Anderson, A.T., 1998. Effects of eruption and lava

constraints on vesiculation. Rev. Miner. 30, 413–445. drainback on the H2O contents of basaltic magmas at Kilauea Stothers, R.B., Wolff, J.A., Self, S., Rampino, M.R., 1986. Basaltic volcano. Bull. Volcanol. 59, 327–344. fissure eruptions, plume heights, and atmospheric aerosols. Geo- Wallace, P.J., Anderson, A.T., 2000. Volatiles in magmas. In: phys. Res. Lett. 8, 725–728. Sigurdsson, H., Houghton, B., McNutt, S.R., Rymer, H., Swanson, D.A., Duffield, W.A., Jackson, D.B., Peterson, D.W., Stix, J. (Eds.), Encyclopedia of Volcanoes. Academic Press, 1979. Chronological narrative of the 1969–71 Mauna Ulu erup- San Diego, CA, pp. 149–170. tion of Kilauea volcano, Hawaii. U. S. Geol. Surv. Prof. Pap. Weill, A., Brandeis, G., Vergniolle, S., Baudin, F., Bilbille, J., 1056, 1–55. Fevre, J.-F., Piron, B., Hill, X., 1992. Acoustic sounder meas- Taitel, Y., Bornea, D., Duckler, A.E., 1980. Modelling flow pattern urements of the vertical velocity of volcanic jets at Stromboli for steady upward gas–liquid flow in vertical tubes. AIChE J. volcano. Geophys. Res. Lett. 19, 2357–2360. 26, 345–354. Williams, S.N., 1983. Plinian airfall deposits of basaltic composi- Thomas, N., Jaupart, J., Vergniolle, S., 1994. On the vesicularity of tion. Geology 11, 211–214. pumice. J. Geophys. Res. 99, 15633–15644. Wilson, L., 1980. Relationships between pressure, volatile content Thorarinsson, S., Sigvaldason, G.E., 1972. The Hekla eruption of and ejecta velocity in three types of volcanic explosion. J. Vol- 1970. Bull. Volcanol. 36, 1–20. canol. Geotherm. Res. 8, 297–313. Thordarson, T., Self, S., 1993. The Laki (Skafta Fires) and Grims- Wilson, L., Head, J.W., 1981. Ascent and eruption of basaltic mag- votn eruptions in 1783–1785. Bull. Volcanol. 55, 233–263. ma on the Earth and Moon. J. Geophys. Res. 86, 2971–3001. Thornber, C.R., 2003. Magma-reservoir processes revealed by geo- Wilson, L., Head, J.W., 1983. A comparison of volcanic eruption chemistry of the Pu’u ‘O’o–Kupaianaha eruption. U. S. Geol. processes on Earth, Moon, Mars, Io and Venus. Nature 302, Surv. Prof. Pap. 1676, 121–136. 663–669. Tilling, R.I., Christiansen, R.L., Duffield, W.A., Endo, E.T., Hol- Wilson, L., Head, J.W., 1994. Mars: review and analysis of volcanic comb, R.T., Koyanagi, R.Y., Peterson, D.W., Unger, J.D., 1987. eruption theory and relationships to observed landforms. Rev. The 1972–1974 Mauna Ulu eruption, Kilauea volcano: an ex- Geophys. 32, 221–263. E.A. Parfitt / Journal of Volcanology and Geothermal Research 134 (2004) 77–107 107

Wilson, L., Sparks, R.S.J., Walker, G.P.L., 1980. Explosive volca- Wolfe, E.W., Neal, C.A., Banks, N.G., Duggan, T.J., 1988. Geo- nic eruptions: IV. The control of magma chamber and conduit logical observations and chronology of eruptive events. U. S. geometry on eruption column behaviour. Geophys. J. R. Astron. Geol. Surv. Prof. Pap. 1463, 1–97. Soc. 63, 117–148. Wolff, J.A., Sumner, J.M., 2000. Lava fountains and their products. Wilson, L., Parfitt, E.A., Head, J.W., 1995. Explosive volcanic In: Sigurdsson, H., Houghton, B., McNutt, S.R., Rymer, H., eruptions: VIII. The role of magma recycling in controlling Stix, J. (Eds.), Encyclopedia of Volcanoes. Academic Press, the behaviour of Hawaiian-style lava fountains. Geophys. J. San Diego, pp. 321–329. Int. 121, 215–225. Woods, A.W., 1993. A model of the plumes above basaltic fissure Wolfe, E.W., Garcia, M.O., Jackson, D.B., Koyanagi, R.Y., Neal, eruptions. Geophys. Res. Lett. 20, 1115–1118. C.A., Okamura, A.T., 1987. The Pu,u O,o eruption of Kilauea Zimanowski, B., Buttner, R., Lorenz, V., Hafele, H.-G., 1997. Frag- volcano, episodes 1–20, January 3, 1983 to June 8, 1984. U. mentation of basaltic melt in the course of explosive volcanism. S. Geol. Surv. Prof. Pap. 1350, 471–508. J. Geophys. Res. 102, 803–814.