Volkanoloji Volkanoloji ( İngilizce ) - 2 85 Chapter 3 Plate VII 1. Ballistic bomb in phreatomagmatic tephra bed surface in the Kissomlyó tuff ring, Hungary. 2. Ballistic bombs damaged buildings during the Usu 2000 eruption, Hokkaido, Japan. 3. Phreatic explosion crater in Usu, Hokkaido, Japan. Note the destroyed house showed on Plate VI, 2. 4. Impact sags in a phreatomagmatic succession of a maar from southern Tenerife. 5. Drop stone-like feature (arrow) caused by a ballistic bomb impacted on a soft and water-saturated accretionary lapilli-rich freshly deposited tephra (a) of the Ság-hegy tuff ring, western Hungary.86 Chapter 3 Plate VIII 1. Plastically deformed impact sag forming twisted and chaotic bed surface of a base surge succession of a maar in southern Tenerife. 2. Hard, strongly diagenised tuff layer (middle in the phreatomagmatic succession of Tihany. Note the V-shaped erosional channels (dashed lines) in cross-section just above the lense cap. 3. Dish structures of a vol- caniclastic succession near Sinker Butte, Idaho. 4. Sand volcano in cross sec- tion from the maar lacustrine units of the Pula maar, Hungary. 5. Clastic dykes intruded into growing lava dome and hyaloclastite unit of Pálháza, Hungary.Volcanic processes are providers of large volumes of volcanic sediment to the surrounding areas of an active volcanic system. They can be classified as either effusive or explosive. During effusive eruptions, various volumes of magma may reach the surface, forming lava flows. Lava flows can provide significant amounts of volcanic debris for secondary, sur- face erosional processes. Such volcanic material can build up large volumes of clastic sediments that may accumulate in the surrounding effusive eruption dominated terrains. Explosive volcanic eruptions, either triggered by magmatic gas exsolution of the emerging magmatic column in a volcanic conduit and/or triggered by magma and water interaction can deposit large volumes of tephra that is ready to be remobilised subsequently by non-volcanic surface processes. In either way, the interpretation of volcanic rocks (both clastic and coherent) need a broader understanding of the relationship between non-volcanic and volcanic successions. Such studies can be performed during detailed investigation of the vol- canic successions, and their relationships to the immediate pre-, syn- and post-volcanic sedimentary successions. This approach is commonly referred to as volcanic facies analysis, and applies similar methods as facies analysis commonly done in non-volcanic sedimentary environments. In this chapter we primarily focus on volcanic sediments, and their importance in understanding the evolution of a volcanic terrain. Volcanic facies Facies is in general a collection of features of a geological unit that define the processes of origin and source as well as the environment of deposition. Facies is in this respect a broad term widely used in geology. Any volcanic facies, per- haps any volcanic rock unit, may define the above basic characteristics. Facies in sedimentology has been defined as the physical, chemical and biologic variations of three-dimensional rock bodies deposited within a specific interval of geo- logical time (PETTIJOHN et al. 1972, 1975).This definition is valid for any sedimentary rock. The definition itself deter- mines that a facies (or volcanic facies) is a well-defined, usually mapable rock unit that bear common textural, compo- sitional, and internal structural features indicative for a well-defined source, transportation and depositional mechanism. Volcanic facies differ from non-volcanic facies in many ways. In order to understand better their formation, the identifi- cation of a single fragmentation history of the juvenile volcanic fragments (in case of primary pyroclastic successions) or multiple fragmentation histories for individual pre- served volcanic clasts (in case of reworked or resedi- mented volcaniclastic rocks) accumulated in a rock unit is essential. Especially in clastic sedi- mentology a useful subdivi- sion of certain facies is used to separate proximal, medial and distal facies (Figure 4.1). 87 Chapter 4 Figure 4.1. Proximal to distal facies variations along an ignimbrite unit (after WRIGTH et al 1981)This subdivision is meaningful since many types of geophysical mass flows (WRIGHT et al. 1981) significantly change their transport and deposition behaviour “en route”. This subdivision is also very useful among volcanic facies, especial- ly in association with pyroclastic density current, laharic or volcanic debris avalanche deposits as well as any secondary, surface erosion resulted volcanogenic mass flow deposit (MATHISEN and MCPHERSON 1991) (Figure 4.2). This type of subdivision in volcanic mass flow deposits is very relative. In base surge deposits accumulated around tuff rings or maars the distance between proximal (Plate I, 1) and distal facies (Plate I, 2) is about 1 km (WATERS and FISHER 1971, CHOUGH and SOHN 1990, VAZQUEZ and ORT 2006). Similar short distance facies variations have been observed at small volume emergent or subaqueous volcanoes such as Surtseyan-style tuff cones (SOHN and CHOUGH 1992, MUELLER et al. 2000, MARTIN 2002). Often particular bedding features (Plate I, 3) are characteristic and restricted to specific distances from source. This has been well documented for the Laacher See Tephra in Germany (FISHER and WATERS 1970, SCHMINCKE et al. 1973). Facies variations may occur over much larger distances in the case of large volume, pyroclastic flow form- ing eruptions. The 1980 eruption of Mount St. Helens reveals facies changes in a distance of a few kilometres from source (CAREY and SIGURDSSON 1982, ROWLEY et al. 1985, FISHER et al. 1987, DRUITT 1992, ALIDIBIROV 1995). Facies changes occur even in longer distances in case of ignimbrite forming events, such as the Campa- nian Ignimbrite in Italy (FISHER et al. 1993). Pro- ximal to distal facies changes can take place over even larger, tens of kilometres in case of large volume ignimbrite erup- tions such as those formed in the Andes (DE SILVA 1989, ORT 1993, LINDSAY et al. 2001). The recogni- tion of proximal to distal facies variations as well as the establishment of a cor- rect stratigraphy of exten- sive ignimbrite sheets can be problematic. Here geochemical studies can prove useful to demonstrate facies changes and to identify facies relationships. This problem is especially pronounced in ancient settings, where the three-dimensional architecture of the otherwise extensive volcanic rock unit is often poorly constrained. This is the case in the Miocene “Rhyolite Tuff Formations” in the Pannonian Basin (CAPACCIONI et al. 1995, PÓKA et al. 1998, SZAKÁCS et al. 1998, LUKACS et al. 2004). In general, in volcanic environments proximal volcanic facies usually include the source volcano (e.g. the volcanic rock units accumulated and formed near to the source vent). In young volcanic terrains, where the original volcanic land- form is more or less still intact, the association of proximal to distal volcanic facies to the geometrical position of the pre- served rocks units in comparison to the source volcano is generally straight forward. In old volcanic successions howev- er, the source volcano may be completely vanished, or displaced by tectonic processes into far position, hindering the identification of proximal to distal facies relationships (Plate I, 4). One of the most important facts in studying old vol- canic successions is, that in most of the cases volcanic rocks accumulated and preserved in basinal settings. Basin fills can be intercalated with distal facies portions of primary to secondary volcaniclastic successions, and only inference of the texture, extension, and geometry of the proximal volcanic deposits is possible (SMITH et al. 1988, BESLY and COLLINSON 1991, MONTANARI et al. 1994, DOSTAL et al. 2003, KATAOKA 2005). The study of basin filling volcaniclastic successions can tell us important details of the deposition environment (e.g. subaqueous versus subaerial) (BARCAT et al. 1989, BULL and CAS 1989, TURBEVILLE 1991, YOGODZINSKI et al. 1996, SCHMINCKE et al. 1997, JERRAM et al. 1999, SCHNEIDER et al. 2001, BREITKREUZ et al. 2002, MARTIN and NÉMETH 2005). In the geological record volcanic sandstones often form the only remnants (Plate I, 5) preserved in the continental sedimentary record (UFNAR et al. 1995, MILLWARD et al. 2000, ZIMMERMANN and BAHLBURG 2003). There is also significant facies change from primary and secondary volcanic successions. In distal regions, primary pyroclastic successions commonly transform laterally into secondary volcaniclastic successions. Such changes often 88 Figure 4.2. Volcanic facies relationships along an active rift zone, complicated by intra-rift faulting and sub-basin development (after MATHISEN and MCPHERSON 1991: p. 30, fig. 5)occur in continental to marine deposit transitions, as many primary pyroclastic density currents enter to lakes or the sea (Plate I, 6) (BULL and CAS 1989, MUELLER 1991, DE RITA et al. 2002). Therefore primary pyroclastic successions, com- monly inter-finger with marine deltaic, turbidity current or deep marine successions (COLE and STANLEY 1994, COUSINEAU 1994). The recognition of such facies changes and the establishment of a three-dimensional relationship between individual rock units may help to establish a basin-wide environmental reconstruction (KANO 1991, MCPHIE and ALLEN 1992, MACDONALD and BARR 1993, FERGUSSON et al. 1994, MANGANO and BUATOIS 1997, NAGY et al. 1999, HATHWAY and KELLEY 2000, DOSTAL et al. 2003, FRANZESE et al. 2003). Volcanic eruptions, especially major arc related strato-volcanic eruptions, immediately influence the surrounding sed- imentary environment, and can provide significant volume of volcaniclastic material into the sedimentary basin. Such composite and long-lived volcanic systems simultaneously provide primary pyroclastic deposits in syn-eruptive time and their reworked derivates during inter-eruptive periods. In this way, a complex primary and secondary volcanic sedimenta- ry record can accumulate, commonly inter-fingered with non-volcanic deposits. It may be further complicated by the fact that horizontally moving volcanic gravity currents (either primary or secondary) obey to physical laws to deposit the mate- rial they transport. The transformations of the physical properties of volcanic gravity currents are the major cause of for- mation of changes in the facies of the accumulating volcanic deposits (FISHER 1983). Such changes can be caused by (1) density separations within the passing particle-laden gravity current due to gravitational settling of particles (gravity trans- formation), (2) velocity variations caused by slope angle changes without significant variations of interstitial fluid content of the body of the gravity current (body transformation), and (3) separations of particles by the turbulent mixing within the boundary between the ambient fluid and the flow surface (surface transformation) (FISHER 1983) (Plate II, 1). Volcanic facies analysis With volcanic facies analyses volcanic deposits (both coherent and fragmented) are studied in an aim to establish hor- izontal and vertical facies relationships that could help to understand the volcanic eruptions, the subsequent reworking processes and the relationship as well as the interaction between volcanic and non-volcanic sedimentary environments. During volcanic facies analysis we can identify general trends of the texture, composition, lateral distribution, magma frag- mentation and eruption style generated the studied volcanic successions. In general, lithofacies studies focus on distinguishable volcanic units, that can be mapped and therefore limit in the field (e.g. in outcrop scale). To define a lithofacies, common, and ready to recognized features such as struc- ture, texture, internal organisation, geometry, and componentry should be identified in outcrop scale. A volcanic lithofacies can contain more than a single bed, and is not necessary the result of a well distinguished volcanic event (e.g. a single pyroclastic density current deposited succession). An identified lithofacies often reveals common vertical and lateral facies varia- tions that can be interpreted as changes in the depositional behaviour of the parental flow. To describe a volcanic lithofacies, similar methods are used as in clas- tic sedimentology. To distinguish certain lithofacies, commonly the grain size characteristics of the deposit (rock) unit is one of the key parameters, since grain size pattern of a certain deposit (rock) has direct link to the physical properties of the transporting and depositing currents (Figure 4.3). The size grading (graded, non-graded, normal or reverse graded, den- sity graded) texture of the unit is an additional useful parameter to distin- guish certain lithofacies. In addition, the internal stratification and clast orientation (e.g. imbrication) are important parameters to define a certain lithofacies (SOHN and CHOUGH 1989, CHOUGH and SOHN 1990). On the basis of these three major parameters, a table-like matrix easily can be cre- ated, and every well-defined rock unit (e.g. bed or group of beds) can be assigned to one of the defined category. The categories may vary (Figure 4.4) in accordance of the usual lithofacies categories used in previous stud- ies and perhaps for practical reasons (e.g. not to be lost in a too detailed descriptive study) (SOHN and CHOUGH 1989, CHOUGH and SOHN 1990, SOHN and CHOUGH 1992, SOHN 1995, NÉMETH et al. 1999, NÉMETH et al. 2001, NÉMETH and WHITE 2003). In a well-defined matrix of rock unit parameters may give certain categories cannot be identified in the field. 89 Figure 4.3. Lithofacies classification diagram used in SOHN and CHOUGH (1989) and CHOUGH and SOHN (1990) to describe volcanic units from the phreatomagmatic volcanoes of the Jeju Island, Korea. In the horizontal “axis” the grain size (from tuff breccia to tuff) and in the vertical “axis” the main bedding features have been used to identify key lithofacies categories90 Figure 4.4. Lithofacies classification diagram used in NÉMETH et al. (2001) to describe volcanic units from the phreatomagmatic volcanoes of the Tihany Peninsula, Hungary. The main categories to identify certain lithofacies were similar to the SOHN and CHOUGH (1989) and CHOUGH and SOHN (1990) lithofacies scheme. Next to bedding characteristics, the dominance of accidental lithic fragments as well as scoriaceous parti- cles in the sediment has been used to distinguish between phreatomagmatic and magmatic explosive erup- tion related deposits Figure 4.5. Lithofacies classification diagram used in NÉMETH and WHITE (2003) to describe volcanic units from the phreatomagmatic volcanoes of the Waipiata Volcanic Field, Otago, New Zealand. The main categories to identify certain lithofacies were similar to the SOHN and CHOUGH (1989) and CHOUGH and SOHN (1990) lithofacies scheme. In this classification scheme the volume of accidental lithic fragments versus juvenile fragments in the deposits has been used to refine the classification schemeAfter describing and interpreting each rock units in this way, a very detailed lithofacies group can be established. In the field perhaps certain cycles in vertical sections and certain vertical facies relationships can be traced. Such cycles and/or common vertical facies relationships can give important information to the transportation and depositing gravity currents, and their physical properties. Certain vertical facies relationships could be identified in specific geometrical distributions, and therefore a further horizontal facies relationship can be established. Such horizontal facies relationships are character- istic for the proximal, medial or distal section of the preserved volcanic succession. As an outcome of such a detailed study, a three-dimensional reconstruction of volcanic processes, their environmental variations and the identification of source regions can be done. In volcanic successions, fragmentation history of juvenile clasts or the relative ratio between juvenile and accidental lithic fragments can also be added as a dividing category (Figure 4.5). In this respect, magmatic and phreatomagmatic fragmentation history of the juvenile clast population of an otherwise similarly transported pyroclastic system can result in two distinct lithofacies (in case other sedimentological parameters are the same). Using assigned facies to describe outcrop-scale volcanic successions can be illustrated in a stratigraphic column 91 Figure 4.6. A graphic stratigraphic log of the phreatomagmatic successions of Tihany, Hungary, detailed with the identified lithofacies categories (NÉMETH et al. 2001)(Figures 4.6 and 4.7). In volcanic field-wide mapping, such columns can help to establish vertical and lateral facies rela- tionships that may be used to reconstruct the volcanic eruptive environment. Differences of volcaniclastic sediments from non-volcanic sediments From a methodological point of view studies and interpretations of volcaniclastic rocks/sediments do not differ from those performed on clastic non-volcanic sediments. However, there are characteristic differences that made volcaniclas- tic studies commonly complicated. Such differences can be taken in account during defining our lithofacies descriptive categories. The specific features of volcanic processes should be treated with care. Descriptive methods should not be mechanically applied during the identification of certain lithofacies. Time-scale of deposition One of the most difficult issues in studies of volcaniclastic sediments/rocks is to understand the time-scale of depo- sition. Usually a bedding plane in normal sedimentary environments clearly demonstrates a time break in sedimentation 92 Figure 4.7. A graphic stratigraphic log of the phreatomagmatic successions of one of thephreatomagmatic volcano of the Waipiata Volcanic Field in Otago, New Zealand (NÉMETH and WHITE 2003)that allows forming a layer physically different to those located below that surface. Such differences may be generated by slight changes of physical, chemical or compositional changes of the depositing sediments subsequently accumulat- ing over a freshly developed bedding plane. But in fact, the time break is in most of the cases significant. In accumulat- ing tephra however, characteristic bedding surfaces can develop without significant time breaks between certain deposi- tional events. Such changes commonly manifest as bed planes and instead of giving evidence of time breaks during the sedimentation giving signs of changes of eruption conditions. In this respect a primary volcaniclastic succession may consists of many well-defined beds but none of them reflect breaks, rejuvenation or erosional events may affect the accu- mulating tephra. Based on direct observation of volcanic eruptions it is well known that characteristic changes could take place in the development of the eruption cloud over the duration of the eruption (Plate II, 2). In Plinian eruptions, the eruption continuously forms a large umbrella like eruption cloud. This eruption plume is initially supported by the upris- ing gas thrust and can reach few tens of kilometres above the volcano. However, there is a fine balance, when the erup- tion column will be charged by expelled pyroclasts that become heavy and the gravitational forces eventually overcome the upward directed forces. In such situations, the eruption cloud will start to partially collapse and the particles fall back to the surface and initiate horizontally moving pyroclastic density currents. In this simple eruption scenario, the accumu- lating tephra will reflect these changes upsection its stratigraphy (Plate II, 3). In spite of the dramatic changes of the transportation and therefore the depositional sequences of the accumulating tephra, the entire pyroclastic succession pro- duced in such events can be volumeneous, and may reflect not more then few hours history, with no significant time breaks. However, the sequence will be strikingly complex, with many well-defined beds in the commonly few tens of metres thick successions. The depositional time of primary volcaniclastic deposits are generally also significantly shorter than that of any other non-volcanic deposits. In extreme case, during magma–water interaction triggered phreatomagmatic explosive eruptions or during scoria cone forming events, few tens of metres thick pyroclastic sequence can accumulate in just few hours (Plate II, 4). In spite of the very short accumulation time frame of such events, the resulting deposits, again, can show complex bedding features, but almost none of them reflecting time breaks of the eruption itself. The complexity of such quickly accumulated sequences clearly associated with the complexity of the eruption dynamics of phreatomagmatic explosive eruptions. Even volcanic events, that are able to mobilise few km 3 volume of material in a single event, such as a formation of volcanic debris avalanches due to volcano collapses, are believed to form within a few days time. The identified deposits still can be very complex, and their studies need a very different approach to studies of other non-volcanic sediments. Volume and aerial extension of deposits The accumulated volume of erupted pyroclastic successions can also differ significantly to other non-volcanic sedi- ments. Extensive ignimbrite sheets for instance can be traced over hundreds of kilometres distances, and their thickness may change insignificantly over large distances. Especially large caldera systems can be associated with such large-vol- ume, commonly very monotoneous (at least from textural point of view) ignimbrite successions. The erupted volume is even more dramatic in sheet-like ignimbrite fields like those formed in South America, which lack apparent known sources. The commonly large extension and volume of volcaniclastic deposits perhaps bear some internal textural vari- ations. Such variations are commonly not obvious in short distances, but can be very pronounced over large (tens of kilo- metres) distances. Textural changes therefore can be very gradual, and only detailed facies analysis can reveal them. In other hand pyroclastic deposits can also be very limited in extension and volume, and could be interbedded with deposits an order larger in volume and lateral extension. Especially phreatomagmatic pyroclastic deposits related to tuff rings and maars are this type of deposits. Lateral textural changes in this type of deposits can be dramatic in very short distances (few hundreds of metres). Cross sectional variations Many volcanic eruptions produce horizontally moving pyroclastic density currents. Such currents can move away from an erupting vent, and gradually depositing tephra “en route” following the changing physical properties of the cur- rents. However, the physical properties of pyroclastic density currents do not exclusively vary in downstream direction but also in lateral extent. Therefore different depositional behaviour characterises their marginal and central zones, which can be identified by significant textural, bedding, and even compositional differences along the cross section of pyroclas- tic flow deposits. Pyroclastic density currents often shift their main depositional axis over time. Such shifts can accumu- late valley and overbank facies over each other creating characteristic vertical facies relationships. In a similar way, other volcaniclastic mass flow deposits (e.g. any laharic deposits) can also be associated with characteristic marginal and cen- tral facies. In this respect identification of cross sectional variations of volcaniclastic mass flows is important to distin- guish from other textural variations caused by longitudinal changes of deposition behaviour. 93Variations in composition Eruption conditions can change dramatically over short times. A sudden entry of water to an erupting vent can switch magmatic gas exsolution driven magma fragmentation to more magma–water interaction driven phreatomagmatic frag- mentation (BARBERI et al. 1989). As a consequence the resulting eruption column will be charged in excavated country rocks that may deposit in well-distinguished beds of fall or pyroclastic density currents. Such changes can produce a vol- caniclastic sequence which maybe deposited from the same eruption cloud or pyroclastic density current but transport- ed different type of clasts. This type of changes can also be gradual, with no apparent sharp marks may result distinguish- able, or easy to recognize beds. Also, the near surface magmatic feeding system of a volcano can change vastly during the eruption. The volcanic conduit can collapse or become wider, changing the style of an ongoing eruption. Typical acci- dental lithic clast horizons may reflect vent clearing events (Plate II, 5). The level of vesiculation of erupting magma can change resulting pyroclasts with different vesicularity. A result of this process is commonly recognised in scoria cone sequences (HOUGHTON et al. 1999). Chemical zonation of a magma chamber also can result in gradual changes of the composition of juvenile fragments in an otherwise texturally uniform pyroclastic succession. Such changes commonly have no obvious manifestation in the accumulated pyroclastic successions, and only detailed geochemical study of the juvenile fragments can reveal its exis- tence (BRIGGS et al. 1993, WILSON et al. 1995, BERESFORD et al. 2000, EDGAR et al. 2002, MAUGHAN et al. 2002, SUMNER and BRANNEY 2002). In other cases chemical zonation of a pyroclastic succession can also be abrupt as the result of a gradual emptying of a layered magma chamber. Thermal properties of pyroclasts Juvenile pyroclastic particles erupted from volcanic eruptions are commonly hot, and their hot temperature can be retained long during transport. Gradual cooling of the depositing pyroclastic system can effect the behaviour of the fresh- ly deposited tephra. In extreme case, ignimbrites in near vent position tend to retain their heat long enough to form weld- ed zones. In more distal position, ignimbrite units can be still hot enough to vapourise entrapped water which may result in secondary explosion craters on the deposit surface (MOYER and SWANSON 1987). Such active ther- mal circulation can create textural features such as hard layers, degassing pipes, or hydrothermal mineralization. To distinguish these features from primary depositional features can be challenging, and detailed facies analyses of the pre- served volcanic succession is neces- sary. Vertical and lateral facies rela- tionships During volcanic facies analysis, certain trends can be outlined both vertical and longitudinal manner. Vertical facies relationships are important to understand the time- scale, apparent shifts of the vent location, or changes in eruption style during the course of the erup- tion (Figure 4.8). Meanwhile the longitudinal facies relationships can tell us important facts about the evo- lution of a volcanic gravity current from the proximal to the distal areas (Figure 4.9). Such changes are com- 94 Figure 4.8. Identified vertical facies variations in relationship to base surge evolution (SOHN and CHOUGH 1989; CHOUGH and SOHN 1990; SOHN 1996). A) Graphic representa- tions of dominant types of vertical facies relationships identified from phreatomagmatic volcanoes of the Jeju Island, Korea. B) 3 main vertical facies relationship from phreatomagmatic volcanoes of the Jeju Island, Korea. Compare the facies codes to the facies table on Figure 4.3. C) A model to associate the identified vertical and horizontal facies variations of the phreatomagmatic volcanoes of the Jeju Island, Korea with the evo- lution of a base surge current, which gradually become more turbulent due to loss of par- ticle concentration “en route”monly related to the flow transformation of certain volcanic gravity currents. During careful facies analyses, very fine changes and relationships can be established as it has been demonstrated from small-volume mafic explosive volcanoes from Jeju Island, Korea (SOHN and CHOUGH 1989, CHOUGH and SOHN 1990, SOHN and CHOUGH 1992, 1993, SOHN 1995, 1996), Sonora, Mexico (WOHLETZ and SHERIDAN 1979, 1983b) or from Linosa, Italy (LAJOIE et al. 1992). However, recognizing facies relationships and connect to them to real natural processes is not easy and ongoing subject of recurring researches of similar fields. Lateral facies variations of tuff ring sequences has been established on the basis of tuff rings and maars of Sonora (WOHLETZ and SHERIDAN 1979, WOHLETZ and SHERIDAN 1983a). Similar studies from Linosa, Italy (LAJOIE et al. 1992) revealed a facies change trend from massive to paral- lel bedded and then dune bedded facies groups larger distances from the source (Figure 4.10). An even more detailed study from Korea (SOHN and CHOUGH 1989, CHOUGH and SOHN 1990), suggested far more complex facies relationships (Figure 4.11) highlighting the role of the level of details how a facies study can be performed in the resulting interpretations. Similar problems are even more evident in larger volume volcanic systems, where outcrop availability, preservation and the large lat- eral extension can hinder the identification of key facies relationships. 95 Figure 4.9. Identified lateral (horizontal) facies relationships in base surge deposits of phreatomagmatic volca- noes of the Jeju Island (SOHN and CHOUGH 1989; CHOUGH and SOHN 1990). Compare the facies codes with the facies table on Figure 4.3 Figure 4.10. Horizontal facies relationships of identified facies in relationship with increasing distance from the source in a tuff ring sequence of Linosa Island, Italy (LAJOIE et al. 1992). Massive facies gradually give places to cross-bedded facies in increasing distance from the source Figure 4.11. Horizontal facies relationships of identified facies in relationship with increasing distance from the source in a tuff ring sequence of Jeju Island, South Korea (SOHN and CHOUGH 1989; CHOUGH and SOHN 1990). Stratified facies gradually give places to undulatory and dune-bedded facies in increasing distance from the sourceVolcaniclastic aprons Large-scale explosive volcanic eruptions spread large volumes of volcaniclastic materials to volcanic fan, alluvial plain, lacustrine and marine settings. The volcaniclastic sediments may become redistributed shortly after an eruption by various reworking processes, such as debris flows, hyperconcentrated flows, dilute fluvial flood flows, turbidity and storm currents (SMITH 1987, SCOTT 1988). Resedimentation of volcaniclastic debris in terrestrial settings is commonly associated with major and catastrophic environmental changes caused by the supply of high sediment load into the basin (PALMER and WALTON 1990, MATHISEN and MCPHERSON 1991, SMITH and LOWE 1991). Large-scale eruptions can poten- tially result in catastrophic impacts over very wide areas and in far-distant places. However, resedimentation induced by large-volume (i.e. hundreds of km 3 ) ignimbrite-emplacing eruptions has not been observed directly and re- mains poorly understood. Volcanism in subaerial and subma- rine settings can accumulate large vol- ume of apron-like deposits. Pyroclastic fans are common around large fre- quently active volcanoes, and could build up thousands of metres thick units. Volcano Mayon in the Phillip- pines (RODOLFO and ARGUDEN 1991), is a well-known example for extensive volcaniclastic fans with interfingered primary and reworked depositional units (Figure 4.12). These types of volcaniclastic fans are commonly named block and ash fans (SIEBE et al. 1993) and can reveal a huge complexity. Large volcaniclastic fans can also develop in alluvial settings, especially where active tectonism and basin subsidence can enhance the inflow of large volumes of volcaniclas- tic sediments from nearby stratocones. Such settings are very common in the geological record and their potential for hydrocarbon exploration is still not fully explored (MATHISEN and MCPHERSON 1991). ALIDIBIROV, M. A. 1995: A model for the mechanism of the May 18, 1980 Mount St-Helens Blast. — Journal of Volcanology and Geothermal Research 66 (1–4), pp. 217–225. BARBERI, F., CIONI, R., ROSI, M., SANTACROCE, R., SBRANA, A. and VECCI, R. 1989: Magmatic and phreatomagmatic phases in explo- sive eruptions of Vesuvius as deduced by grain-size and component analysis of the pyroclastic deposits. — Journal of Volcanology and Geothermal Research 38 (3–4), pp. 287–307. BARCAT, P. F., CORTINAS, J., NEVISTIC, V. and ZUCHI, H. 1989: The Huriwai braidplain delta of New Zealand; a late Jurrasic, coarse- grained, volcanic-fed depositional system in a Gondwana forearc basin. — In: NEMEC, W. and STEEL, R. J. (Eds): Fan deltas: sed- imentology and tectonic setting, Blackie & Son, pp. 431–444. BERESFORD, S. W., COLE, J. W. and WEA VER, S. D. 2000: Weak chemical and mineralogical zonation in the Kaingaroa Ignimbrite, Taupo Volcanic Zone, New Zealand. — New Zealand Journal of Geology and Geophysics 43 (3), pp. 639–650. BESLY, B. M. and COLLINSON, J. D. 1991: Volcanic and tectonic controls of lacustrine and alluvial sedimentation in the Stephanian coal-bearing sequence of the Malpas-Sort Basin, Catalonian Pyrenees. — Sedimentology 38, pp. 3–26. BREITKREUZ, C., CORTESOGNO, L. and GAGGERO, L. 2002: Crystal-rich mass flow deposits related to the eruption of a sublacustrine silicic cryptodome (Early Permian Collio Basin, Italian Alps). — Journal of Volcanology and Geothermal Research 114 (3–4), pp. 373–390. BRIGGS, R. M., GIFFORD, M. G., MOYLE, A. R., TAYLOR, S. R., NORMAN, M. D., HOUGHTON, B. F. and WILSON, C. J. N. 1993: Geochemical zoning and eruptive mixing in ignimbrites from Mangakino Volcano, Taupo Volcanic Zone, New-Zealand. — Journal of Volcanology and Geothermal Research 56 (3), pp. 175–203. BULL, S. W and CAS, R. A. F. 1989: Volcanic influences in a storm- and tide-dominated shallow marine depositional system: The Late Permian Broughton Formation, southern Sydney Basin, Kiama, NSW. — Australian Journal of Earth Science 36, pp. 569–584. CAPACCIONI, B., CORADOSSI, N., HARANGI, R., HARANGI, S., KARATSON, D., SAROCCHI, D. and VALENTINI, L. 1995: Early Miocene pyro- clastic rocks of the Bukkalja ignimbrite field (North Hungary); a preliminary stratigraphic report. — Acta Vulcanologica 7 (2), pp. 119–124. CAREY, S. N. and SIGURDSSON, H. 1982: Influence of particle aggregation on deposition of distal tephra from the May 18, 1980, erup- tion of Mount St-Helens Volcano. — Journal of Geophysical Research 87 (NB8), pp. 7061–7072. CHOUGH, S. K. and SOHN, Y. K. 1990: Depositional mechanics and sequences of base surges, Songaksan tuff ring, Cheju Island, Korea. — Sedimentology 37 (6), pp. 1115–1135. 96 Figure 4.12. Faices model of the Mayon lahar system in Philippines (after RODOLFO and ARGUDEN 1991). 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ZIMMERMANN, U. and BAHLBURG, H. 2003: Provenance analysis and tectonic setting of the Ordovician clastic deposits in the southern Puna Basin, NW Argentina. — Sedimentology 50 (6), pp. 1079–1104. 99100 Chapter 4 Location map 1 — Laacher See, Germany 2 — Mount St. Helens, Washington, USA 3 — Campania, Italy 4 — Pannonian Basin, Central Europe 5 — Jeju Island, Korea 6 — Sonora, Mexico 7 — Linosa Island, Italy 8 — Volcan Mayon, Phillippines 9 — Caldera del rey, Tenerife, Spain 10 — Dolomites, Italy 11 — Coastal Coordillera, Chile 12 — Volcan Villarica, Chile 13 — Waipiata Volcanic Field, New Zealand 14 — Ambrym Island (Marum volcano) Vanuatu 15 — Altiplano, Chile 16 — Tihany Peninsula, Hungary101 Chapter 4 Plate I 1. Proximal base surge facies from the crater rim of Crater Elegante maar in Sonora, Mexico. 2. Distal base surge facies about 1500 m away of the crater rim of Planchon, Mendoza, Argentina-Chile border. 3. Dune-bedded base surge beds from the distal section of the phreatomagmatic succession along the Snake River, Idaho, USA. 4. Triassic volcaniclastic siltstone (Pietre Verde) bed (greenish bed) interbedded with normal siliciclastic turbidite beds from the Dolomites, northern Italy. 5. Volcaniclastic sandstone successions from the Jurassic La Negra Formation, Northern Chile. 6. Pumiceous deposits from beach setting in southern Tenerife. A) Dewatering pipes in the pumiceous deposits formed when ign- imbrite entered to the shallow sea, beach environment, and entrapped water that find escape path to the surface after deposition. B) Cauliflower shape pumiceous lapilli as evidence of sudden contact with cold sea water upon the ignimbrite entered to the sea.102 Chapter 4 Plate II 1. Transitional deposits of a block and ash fan of the Calbuco volcano, southern Chile. A) Overview of the thickly bedded, coarse grained volcaniclastic succession, deposited by debris flows from lahars B) Charcoal fragments in pyroclastic flow deposits initiated laharic succession of the Calbuco volcano, southern Chile. 2. Sub-Plinian to Vulcanian eruption cloud over Marum volcanic cone complex in Ambrym Island, Vanuatu. Note the large drifted eruption cloud with well-defined zones of slightly higher cloud. These zones formed in previous Vulcanian explosive eruptions, but the cloud was drifted quickly away. Such pulsating style of eruption can produce characteristic cyclic deposits. 3. Plinian-type of eruptions are associated with initial Plinian-type fall deposits, that can accumulate clast supported, pumice-rich, rel- atively well-sorted plinian fall deposits such the case in this section near to Rotorua, North Island, New Zealand. 4. Small volume monogenetic volcanoes such as the scoria cones and tuff rings of the Auckland Volcanic Field in New Zealand can form in few hours, and can accumulate tens of metres thick pyroclastic successions. 5. Vent and volcanic conduit dynamics can be closely followed in any pyroclastic succession such as the phreatomagmatic units of the Pula maar in Hungary. Beds richer in accidental lithic fragments represent volcanic conduit collapse and/or vent clearing events in the course of the relatively short-lived volcanic eruption.Small-volume volcanic eruptions are commonly referred to be monogenetic form tephra cones, rings, or mounds con- sisting of bedded pyroclastic deposits that are composed by fallout, density currents and/or down slope remobilisation of tephra (CONNOR and CONWAY 2000, VESPERMANN and SCHMINCKE 2000). Monogenetic volcanic fields are commonly related to explosive eruptions (phreatomagmatism) driven by magma–water interaction associated with magma interacting with shallow or deep ground water and/or surface water sources (WHITE 1991a). In many cases seasonal climatic changes as well as the ratio of available surface and ground water play an important role in the formation of different types of volcanic landforms (ARANDA-GOMEZ and LUHR 1996, CARN 2000, NÉMETH et al. 2001, SIEBE et al. 2005). From these reasons a great variety of volcanic landforms are expect- ed to develop especially in low lands where the hydrogeology of the country rocks maybe very complex (WHITE 1991a, 1991b). The resulting volcanic landform especially in such settings is strongly depending on the nature of the pre-eruptive surface (e.g. depositional environment), lithology and mechanical properties of volcanic conduit wall rocks, vent geom- etry, type and availability of external water (LORENZ 1987). Most volcanological studies on monogenetic volcanic land- forms are based on young volcanoes and focus on their short term morphological changes, tephra transport and deposi- tional processes in syn- and post-eruptive times. Due to the erosion the exposed inner architecture of monogenetic vol- canoes reveal volcanic litho-facies that bear important information on the eruptive mechanism of the volcano (LORENZ and KURSZLAUKIS 2007). Monogenetic volcanic fields commonly consist of large numbers of volcanic clusters and/or alignments that may reach the amount of hundreds of volcanoes in a volcanic field (CONNOR 1987, 1990, CONNOR et al. 1992, CONDIT and CONNOR 1996, CONVAY et al. 1998, CONNOR and CONVAY 2000, CONNOR et al. 2000, VALENTINE et al. 2006). Such large volcano numbers in a volcanic field may evolve over millions of years. Over such long time periods (thousands to mil- lions of years) individual volcanoes may erode significantly, and a volcanic field maybe a group of variable eroded vol- canic landforms preserved on a gradually eroding landscape (WHITE 1991b, KONECNY et al. 1999, KONECNY and LEXA 2000, NÉMETH and WHITE 2003b, 2003a, LORENZ and HANEKE 2004). Over longer time periods, a relatively uniform landscape could be dissected, lowered and commonly inverted having preserved clusters of formerly low-lying zones of the syn-eruptive landscapes in elevated positions (NÉMETH and MARTIN 1999). In this point of view, monogenetic volcanic fields are very useful to characterize the course of erosion of the sur- rounding syn-eruptive landscape over the time length during which the volcanic field was active and after the cessation of the last eruption. Especially volcanic erosion remnants of volcanoes in a same age cluster may help to draw a geomor- phic horizon for the same time horizon (NÉMETH and MARTIN 1999). With substantial enough age clusters of volcanic erosion remnants and good enough age data for such clusters a refined erosion history of an often large (hundreds of km 2 ) area can be reconstructed. In this geomorphic reconstruction a combination of general geomorphic consideration of ero- sion trends of certain volcanic landform such as scoria cones, tuff rings, tuff cones, maars are critical. There are many relatively mechanical ways to re-establish the erosional stage of a certain monogenetic landform and therefore estimate the erosion stage of the syn-eruptive landscape on which such volcano erupted. Especially in older (millions of years) volcanic fields, erosional remnants of phreatomagmatic volcanoes bear important information on the syn-eruptive coun- try rock stratigraphy. 103 Chapter 5Monogenetic volcanic fields Volcanic activity in terrestrial settings often results in the formation of volcanic fields rather than single volcanic edi- fices (CONNOR and CONVAY 2000, WALKER 2000). Volcanic fields, especially basaltic ones, are common volcanic sys- tems on Earth (WALKER 1993). They can develop as a cluster of small-volume volcanoes such as the Hopi Butte (Plate I, 1) in Arizona (WHITE 1991b), or around a central volcano, such as the volcanic field around Lamongan (Figure 5.1), in Java (CARN 2000). Monogenetic volcanic fields are those in which individual volcanoes (mainly basaltic) commonly form during single episodes of volcanic activity, without subsequent eruptions, while the volcanic field as a whole may be active for millions of years (WALKER 1993). Intracontinental volcanic fields commonly are characterized by low magma supply rates over relatively long periods of time (millions of years) (TAKADA 1994, CONNOR et al. 2000). They typically consist of scattered volcanic vents that are often considered to be monogenetic as they apparently never made it into the construction stage of stratovolcanism (WALKER 1993). In fact, these volcanoes are generally small in size and in volume of accumulated eruptive products, but on closer inspection they do often show signs of multiple eruption his- tories, and therefore their architecture can be complex regardless of their small size. It is also notable, that volcanic fields in continental settings are often associated with large shield volcanoes and lava flow fields (GREELEY 1982, WALKER 1993, HASENAKA 1994, NÉMETH 2004). Phreatomagmatic volcanoes in a volcanic field are commonly associated with low-lands or valleys (LORENZ 1973, 1986, LORENZ and BÜCHEL 1980). Magmatic explosive eruptive centres and extensive lava fields are commonly located in elevated lands or areas with limited water availability. Lava flows are commonly confined to valleys or stopped behind syn-volcanic geomorphic barriers. Lava flows may not leave their source vent zone, forming lava lakes, or filling wide craters of phreatomagmatic volcanoes. Distribution of different styles of vents gives vital information of the syn-volcanic landscape drainage system as well as its physiography (LORENZ and BÜCHEL 1980). Identification of widespread 104 Figure 5.1. Volcanic field around the Lamongan stratovolcano in East-Java (after CARN 2000: p. 84)phreatomagmatism in many fields suggests extensive surface and ground water availability of the region in syn-volcanic times. Volcanic fields in terrestrial settings, especially those developed in a fluvio-lacustrine basin such as the western Pannonian Basin Mio/Pliocene fields (MARTIN and NÉMETH 2004) or the Snake River Plain (GODCHAUX et al. 1992, WOOD and CLEMENS 2004, NÉMETH and WHITE 2007) volcanic fields are great volcanological interest both from vol- canological and palaeogeomorphological point of view. The relatively long volcanic history of such volcanic fields and the adjacent lake\fluviatile environment makes them an ideal area for studying the sublacustrine, peri-lacustrine and post- lacustrine volcanism, the palaeogeomorphological evolution. Such fields have a great opportunity for developing our knowledge about eruption mechanisms resulting from magma–water interactions across the entire magma\water-ratio spectrum with the special relations to the palaeoenvironment, and the related palaeohidrology and physical characteris- tics of the pre-volcanic units. Fundamental physical characteristics of volcanic fields that are the focus of current research include 1) the number, type and eruption history of individual vents (CONNOR 1990, SIEBE et al. 2005, VALENTINE et al. 2006); 2) the timing and recurrence rates of the volcanic eruptions in a given volcanic field (TANAKA et al. 1986, CONDIT and CONNOR 1996, CONVAY et al. 1998), 3) the distribution of vents and volcanic complexes (CONNOR 1987), and 4) the relationship of vol- canic fields and the volcanoes within them to tectonic features such as basins, faults and rift zones (CONNOR et al. 1992, STAMATAKOS et al. 1997, CONNOR et al. 2000). In general there are three major elements to be considered in the ascent and emplacement of magma either on Earth or other planets, and each strongly depends on the physical properties and structure of the lithosphere encountered by the magma. The three factors are (WALKER 1989): 1) magma generation and buoyancy 2) rheological boundaries in the lithosphere and 3) density boundaries in the lithosphere. In addition to these factors, the stress field (local and regional) plays an important role in controlling magma ascent which is generally relat- ed to the structural features of the lithosphere encountered by the magma. Monogenetic volcanoes are traditionally referred as those volcanoes that erupt only once during their eruptive histo- ry. They are small and occur as scoria cones (Plate I, 2), tuff cones (Plate I, 3) and rings (Plate I, 4), and maars (Plate I, 5). They form from typically short-lived single and brief eruptions. Their characteristic feature is that the duration of the eruption is usually shorter than the solidification time required for the feeding system to provide the melt for the erup- tion. This definition maybe useful to classify volcanoes that erupt mafic magmas and produce small-volume cinder cones associated with variable long lava flows. Since phreatomagmatic volcanoes such as maars and tuff rings are considered to be the wet equivalent of scoria cones (LORENZ 1986), a general preconception of the short eruption duration of phreatomagmatic volcanoes is accepted. Eyewitness of a few historic maar volcanic eruptions such as Ukinrek (Alaska) (KIENLE et al. 1980, SELF et al. 1980, ORT et al. 2000) and or Rininahue (Chile) (MÜLLER and VEYL 1956) support this working hypothesis that such eruptions usually last a short time. Calculation of the necessary melt involvement into such eruptions commonly gave a very low proportion of primary magma that needs to create such volcanoes. This conclusion is supported by the componentry analyses of phreatomagmatic tephra in many places in various volcanic settings. To solidify such small volume of melt in the feeding system it may need less time than the a few days or weeks. Conversely, there are reports from scoria cones which show gradual transition toward composite volcanoes, and they are hard to clas- sify in term of monogenetic and polygenetic systems (MCKNIGHT and WILLIAMS 1997). On the basis of historic eruption it is clear, that most of these volcanoes are large in volume and the deposited tephra commonly consists of a great diver- sity of eruptive products from strikingly different fragmentation history of the melt (e.g. magmatic vs. phreatomagmat- ic). Especially phreatomagmatic volcanoes are often associated with scoria cones, spatter cones and lava flows and they form together a volcanic complex. Such volcanic complexes are closely spaced individual volcanoes that individually may fulfill the requirement of the sensu stricto term monogenetic, however, to identify these features in ancient setting maybe problematic if not impossible. Monogenetic volcanoes may form in any type of geological environment, however, their volcanic landform strong- ly depends on the water availability and therefore the environment where they erupt. In fully subaqueous environments either in sub-lacustrine or sub-marine settings lensoidally shaped pyroclastic mounds form (WHITE 1996). These mounds consist of flat lying pyroclastic density current deposits (WHITE 2000). In shallow subaqueous environment after the construction of a pyroclastic mound, eruption clouds and directed pyroclast jets breach the water surface and form steep flanked tephra cones over a basal tephra mounds (BELOUSOV and BELOUSOVA 2001, WHITE 2001). In fully subaerial settings (Figure 5.2) when magma interact with near surface water tables and/or very shallow lake, sea or river water, gently dipping broad tephra rings develop (HEIKEN 1971). The tephra ring edifice consists of alternating base surge and phreatomagmatic fall (Plate I, 6) tephras (VESPERMANN and SCHMINCKE 2000). When magma interact with ground water, hole-in-the-ground, maar volcano forms (Plate I, 7) that is surrounded by flat lying tephra beds (VESPERMANN and SCHMINCKE 2000). The crater floor of maars undercut the syn-eruptive surface. The maars have four distinct part (Figure 5.3) in cross sectional view (LORENZ 1986, WHITE 1991b, LORENZ and KURSZLAUKIS 2007); 1) root zone with intrusions that are often mixed with matrix of conduit filling volcaniclastic debris and collapsed country rock blocks; 2) a lower diatreme which is the deep subsurface zone of the conduit filling primary and intra-vent mix- ing of different origin volcaniclastic debris; 3) an upper diatreme that consists of near surface primary pyroclastic 105deposits and 4) the crater lake setting that is an accumulation of different origin sediments in a small sedimentary basin with variable influence from background sedimentation. The above mentioned zones of a maar volcano maybe exposed in accordance to the level of erosion of the surrounding landscape (Plate II, 1). Therefore the clear identifi- cation and the geometrical consideration of the maar volcanic architecture bear vital information of the estimation of the level of erosion. Using eroded monogenetic volcanic fields to calculate long term erosion and to estimate possi- ble palaeogeography of a region are powerful tools to reconstruct syn-volcanic landscapes. Therefore such works could give vital information of a region‘s landscape evolu- tion. Maar/diatreme volcanoes are often the only sites where already eroded pre-volcanic sediments may have been preserved, thus their study may give information to understand the stratigraphy. Phreatomagmatic explosive phases occur in almost every small-volume volcanoes especially in valley settings with good water availability. Subsequent magmatic explosive eruptions, however, are commonly accompanied with such volcanoes in the wan- ing phase of the eruption history of such volcanoes (ARANDA-GOMEZ et al. 1992). This phase commonly pro- duces large scoria cones accompanied by lava effusion commonly erupting inside the crater (Plate II, 2) (ARANDA- GOMEZ et al. 1992). Phreatomagmatic volcanic fields commonly erupted into fluvio-lacustrine basins, therefore the recognition of great diversity of phreatomagmatic volcanoes in ancient settings almost certainly means that the syn- volcanic landscape is rather basin-like, and/or valley settings. Such an interpretation may alter the extrapolation of the estimated landscape erosion data. It also recently has been highlighted that volcanic activity may rejuvenate from time to time in an otherwise small-volume volcano that has been considered to be monogenetic. This may challenge the „monogenetic“ characteristics of such volcanism. Recognition of such processes in the preserved geological records of such small-volume volcanoes must be taken seriously in the estimation of the landscape erosion of the syn-erup- tive settings. In the next paragraphs we list a few basic considerations that may alter the estimation of the certain syn- volcanic landscape erosion level. Lava fields, spatter cones, and scoria cones Hawaiian to Strombolian-type eruptions build up spatter, scoria and/or cinder cones ((Plate II, 3). Eruption rate, volatile content, magma composition and temperature are the most obvious controlling factors during their eruption (HOUGHTON and SCHMINCKE 1989, HOUGHTON et al. 1999, VESPERMANN and SCHMINCKE 2000). 95% of observed cinder- 106 Figure 5.2. Schematic cross sections of the three major types of monogenetic volcanoes (after CAS and WRIGHT 1988: p. 377, fig. 13.17) Figure 5.3. Cross section of a typical maar (after LORENZ 2007: p. 290, fig. 1)cone eruptions lasted less than a year in contrast to composite volcanoes formed from multiple eruptions over thousands of years — an important notion in view of the hazard assessment. Comparative morphology of scoria cones is a useful dating tool, however, new researches suggest that their erosion could be more complex. Rare basaltic Plinian eruptions are poorly known but dangerous volcanic phenomena. Shield volcanoes (Plate II, 4) are common and give the major sources of lavas in intraplate provinces (WALKER 1993). Eruptions of large Hawaiian-type volcanic centers usually relat- ed fissure-vent systems (Plate II, 5), but in small plains-basalt province eruptions related to central vent systems (JOHNSON 1989), but there are several examples where shield volcanoes developed along basement fissure systems (JOHNSON 1989). The individual lavaflows associated with intracontinental volcanic fields tend to be around 5 to 10 km long however excep- tionally long lava flows are also known (CONNOR and CONVAY 2000, KILBURN 2000). The total thickness of the lava cover could reach several tens of meters and could cover significant portions of a volcanic field. Adjacent to the lava field Strombolian scoria cones and Hawaiian spatter cones are common features in a continental volcanic field. These volca- noes are commonly point sources of lava flows (Figure 5.4). Basaltic volcanic fields commonly accompanied by exten- sive lava fields ranging from aa’ to pahoe- hoe types of lava flows (KILBURN 2000). Lava flow fields bear characteristic sur- face morphological features such as tumuli (Plate II, 6), sky rise, whale back humps, lava tubes, and pressure ridges (KILBURN 2000). Such surface morpho- logical features are commonly large in dimension (tens of metres) and character- istic for eruption rate, composition, pre- flow morphological features and composi- tion of the flow (KILBURN 2000). Among these features, tumuli, are whale-back- shaped uplifts are common in most of the pahoehoe lava flow fields, such as the Deccan, India (DURAISWAMI et al. 2001), Hawaii, USA (WALKER 1991), Etna, Italy (CALVARI et al. 2003), Iceland (MATTSSON and HOSKULDSSON 2005), or in Eastern Australia (OLLIER 1964, WILMOTH and WALKER 1993). Tumuli are positive topographic features with wide range of slope angles and surface morphological features that are common on pahoehoe lava flow fields. Commonly three types of tumuli are distinguished on shield volcanic systems, such as (1) lava-coated tumuli, (2) upper-slope tumuli and (3) flow- lobe tumuli in accordance to their distance from their source (ROSSI and GUDMUNDSSON 1996). Flow-lobe tumuli are com- mon in the medial and distal parts of pahoehoe flow fields, whereas the other two tumulus types are more frequent in the proximal parts of the flow fields (ROSSI and GUDMUNDSSON 1996). The flow- lobe tumuli are significantly larger, have shallower flanks, and do not have extensive outflows from the cracks in comparison to more proximal types of tumulus (ROSSI and GUDMUNDSSON 1996). They commonly group into cluster in gentle slopes of shield volcanoes, such Hawaii (WALKER 1991). Large tumuli are comparable in size to small lava spatter cones, and therefore their recognition from morphological point of view is important to establish the eruption history of volcanic fields,. It has been demonstrated that recognition of tumuli features and their characteristic surface textures may give hint to quantify eruption duration (MATTSSON and HOSKULDSSON 2005). There are thin lava foot breccias between the individual lava beds. 107 Figure 5.4. Interpretation of effusive and explosive eruptive processes responsible for the formation of the Lathrop Wells volcano in the Yucca Mountain, Nevada (VALENTINE et al. 2006: fig. 17)108 Spatter cones consist of a near vent strongly baked, red, slightly bedded sequences with large spindle or highly vesic- ulated fluidal bombs (Plate III, 1). These deposits usually reflect strong reworking of volcaniclastics in near vent posi- tion. Spatter cones and scoria cones can build up steep spatter and agglutinate piles, that can collapse gravitationally (Figure 5.5), or driven away by moving lava flow initiated from the flank of the cone (Plate III, 2). Strombolian scoria and spatter deposits are common in relation with maar volcanic centers. Even maar volcanic centers may produce phases of Hawaiian and Strombolian-style eruptive activity from several distinguished eruption points, leading to agglutinates or even clastogenic lavas. There are examples in the Tihany Volcano maar volcanic Figure 5.5. Model of spatter rampart collapse after SUMNER 1998complex for this kind of deposits in the northern part of the maar complex (Gödrös–Diós) (NÉMETH et al. 1999). A remnant of Strombolian scoria cone in the Füzes-tó region in the Bakony–Balaton Highland Volcanic Field, Hungary shows near vent scoriaceous volcaniclastic breccia in peperitic matrix representing water saturated slurry in the vent during the Strombolian activity (NÉMETH and SZABÓ 1998). Magmatic explosive and/or degassing processes as result of the fragmentation of the uprising mafic magma lead- ing the formation of scoria cones with common welded core zones (VESPERMANN and SCHMINCKE 2000). The textur- al characteristics of the pyroclasts, such as high vesicularity, fluidal shape, and the dark, often red colour (Plate III, 3) indicate a magmatic degassing and fragmentation history due to Strombolian-style explosive eruptions (JAUPART and VERGNIOLLE 1988, VERGNIOLLE and BRANDEIS 1996, VERGNIOLLE et al. 1996, SUMNER 1998, VESPERMANN and SCHMINCKE 2000). The closely packed, slightly oriented texture of lava ash and lapilli rich pyroclastic rocks are inter- preted to be the result of Hawaiian-style lava fountaining (THORDARSON and SELF 1993, VESPERMANN and SCHMINCKE 2000, WOLFF and SUMNER 2000, SUMNER et al. 2005). The common intercalation of scoria beds of scoria cones with welded fall out deposits and/or clastogenic lava flows indicate a sudden and common change in eruption style from Strombolian to Hawaiian and vice versa (Plate III, 4) (PARFITT and WILSON 1995, PARFITT et al. 1995, WILSON et al. 1995). The presence of pyroclastic breccias, lapilli tuff and tuff interbeds in scoria cones are common signs of a phreatomagmatic influence on the eruptions (HOUGHTON and HACKETT 1984, HOUGHTON et al. 1991, DOUBIK and HILL 1999, HOUGHTON et al. 1999). Scoria cones are the most common subaerial volcanic landforms on Earth and are generally considered to be a result of mild explosive eruption of mafic magmas in a short period of time (days, weeks) (VESPERMANN and SCHMINCKE 2000), however long-lived scoria cone eruption such as Parícutin in Mexico was active for 9 years (LUHR and SIMKIN 1993). In spite of the numerous scoria cones associated with volcanic fields and central volcanoes (e.g. along rift zones) there are only a few detailed studies that have been carried out on their architecture (MCGETCHIN et al. 1972, CHOUET et al. 1974, MCGETCHIN et al. 1974, MCGETCHIN and SETTLE 1975, HEAD and WILSON 1989, RIEDEL et al. 2003). Scoria cones inner part usually consists of welded agglutinate (Figure 5.6). Such welded part usually more resistant against erosion and can be preserved long time (Plate III, 5). Detailed analyses of deposits preserved on scoria cones lead to the clarification of the role of the shal- low seated magmatic system in the control of the explosive eruptions of such volcanoes (HOUGHTON et al. 1999). Among the identified parameters the vari- ations in degassing patterns, magma ascent rates and degrees of interaction with external water are thought to be responsible for sudden changes in the eruption sequence from deposits representative for “wet” and “dry” eruption conditions (HOUGHTON et al. 1999). In general scoria cone-forming eruptions are linked to Strombolian-type activity driven by magmatic fragmentation occurring in the near sur- face region of the open volcanic conduit (BLACKBURN and SPARKS 1976, HOUGHTON et al. 1999). Among scoria cones a great variety has been observed and described, which show gradual transi- tions between Hawaiian lava fountaining (Figure 5.7) to moderate Strombolian-type eruptions. It has been suggested that magma ascent speed is the most important factor causing such transitions, with gas content and viscosity also influencing the ascent speed at which the transition occurs (PARFITT and WILSON 1995, PARFITT et al. 1995). A decrease in gas content does not cause a transition from Hawaiian to Strombolian activity, but instead causes a transition to passive effusion of vesicular lava (PARFITT and WILSON 1995). Some authors suggest that a change from Hawaiian to Strombolian-style requires a significant reduction in magma ascent speed (PARFITT and WILSON 1995). Among many of the recently identified deposits related to basaltic explo- sive volcanism Tarawera (New Zealand) 1886 eruption is a classical example of characteristic mafic (sub)-Plinian- style (HOUGHTON et al. 2004). “Violent Strombolian” eruptions are explosive eruptions of mafic magma charac- terised by eruption column heights <10 km, voluminous ash production, and simultaneous lava effusion (HOUGHTON 109 Figure 5.6. Scoria cone structure (after VESPERMAN and SCHMINCKE 2000) 1 — initial phreatomagmatic units, 2 — strombolian units with intercalated phreatomagmatic beds, 3 — strombolian eruption formed cone facies, 4 — post-strombolian talus, 5 — distal fallout tephra, CF = crater facies, UCF = upper crater facies, WF = wall facies, TS = talus slopeet al. 2004). The mechanism of gener- ation, fragmentation, transportation and deposition of ash in these erup- tions is poorly understood, however such eruption was just recently identi- fied to have occurred not only during eruption of composite volcanoes but also in single scoria cones. The best example for such eruptions is Volcán Parícutin, Mexico (1943–1953). Eruptive products of Parícutin erup- tion include thick tephra deposits of alternating ash and lapilli beds few kilometres from the cone building pyroclastic units (LUHR and SIMKIN 1993, NEWTON et al. 2005). Comparative studies indicated that other young monogenetic cones in the Michoacán, Mexico region may have erupted in a similar style as Parícutin (HASENAKA and CARMICHAEL 1985). Further field evidences from the western Transmexican Volcanic Belt (Plate III, 6) indicate that mafic violent Strombolian-style eruptions among scoria cones may be more common than expected (MARTIN and NÉMETH 2006). Maar – diatreme volcanoes Maars are small-volume volcanoes that are the second most common volcanic landforms on Earth (LORENZ 1985 1986). Maar volcanoes have characteristically wide and deep craters, commonly referred as ‘hole-in-the-ground’ features (LORENZ 1985). Magma and water explosive interaction is considered to be the main driving force that causes their creation (LORENZ 1986, ZIMANOWSKI et al. 1991, ZIMANOWSKI et al. 1997, BUTTNER et al. 2002). The sudden formation of steam during magma and water interaction produces a steam explosion that disrupts the country rocks and creates a mass deficit in and around the explo- sion centre, which leads to the formation of a collapse crater. The explosion locus gradually migrates downward (Figure 5.8) due to the gradual exhaust of water sources to fuel magma–water interac- tion in the explosion sites (LORENZ 1986). Maar volcanoes are not a strictly or well characterised group of volcanoes. There is an almost continuous transition to scoria (cinder) cones, which are considered to be the most common volcanic landforms on Earth. Scoria cones have frequently had only a short-lived eruptive histo- ry, when magma– water interaction took place. In this respect, maars may be viewed as ‘wet’ equivalents of scoria cones. Maar volcanoes are complicated, but small volume volcanoes if we con- sider that they commonly form groups or clusters, or have struc- turally-controlled alignments. In volcanic fields, the volcanoes commonly fall between the two end-members such as maars and scoria cones and therefore the study of such landforms should include both. There exist two models on the formation of maar-diatreme volcanoes: the magmatic model and the phreatomagmatic model. The magmatic model is especially concerned with ultra- basic, ultramafic and carbonatitic magmas. It invokes volatile rich fluid magmas which, close to the Earth´s surface, fragment the country rocks thus forming progressively from deep- er levels to almost the surface the irregular shaped root zone. Explosive breakthrough to the surface is supposed to result in the formation of the maar crater and then via downward propagating fluidization of the root zone contents to shaping 110 Figure 5.8. Theoretical model of the evolution of a maar due to downward migration of explosion locus in time (after LORENZ 2007: fig. 2) Figure 5.7. Hawaiian lava fountaing model (after CAREY and BURSIK 2000) of the cone-shaped diatreme, and mixing of the diatreme clasts (WOOLSEY et al. 1975, MCCALLUM et al. 1977, WOLFE 1980, MITCHELL 1986). The phreatomagmatic model, in contrast, invokes explosive interaction of rising magma with ground water, originally close to the surface and then downward penetration of the sites of explosion (LORENZ 1985). The various individual explosion sites or chambers jointly form the root zone. Ejection of explosively fragmented country rocks leads to a mass deficiency and consequently to collapse of the overlying rocks. Via these processes the diatreme forms and, in principle, it represents a collapse feature like a sink hole. Downward explosive penetration of the root zone on its own feeder dyke and consequent collapse phases of the diatreme leads to a growing diatreme and a growing maar crater above (LORENZ 1986). Phreatomagmatic explosions and eruptions represent the key mechanism for the formation of maar-diatreme volcanoes, irrespective of magma composition or host rock type. However, it is environmental conditions (including magma and host-rock characteristics) that apparently cause quite a variability in size and shape of maars and the characteristics of their deposits. Furthermore, phreatomagmatism in maar vol- canoes can occur with eruptive stages of purely magmatic explosivity, mixed magmatic/phreatomagmatic erup- tions, or even final stages of lava lake effusion. Because the formation of maars is driven by the interaction of ground water and uprising melt, after the collapse of the crater, ground water inflow quickly fills the volcanic depression and creates a deep crater lake (BÜCHEL and LORENZ 1993). Maar lakes are steep sided, often surround- ed by unstable steep tephra cliffs that can erode into the lake quickly (Plate IV, 1). Maars produce complex deposits both inside and outside of their craters. The record of the eruption is complicated by complex interactions between magmatic and phreatomagmatic processes in the conduit, transport processes in the vertical and lateral currents, and depositional processes. A diatreme is the substructure of a maar crater and its tephra ring (LORENZ 1986, WHITE 1991b, LORENZ and KURSZLAUKIS 2007). In maar-diatreme volcanoes a large amount of fragmented country rocks and commonly juvenile lapilli and bombs are ejected. Formation of these lapilli and bombs is attributed to thermohydraulic explosions in the root zones of the diatremes. Sections of fragmented country rocks may collapse or slide into the partially evacuated root zone and thus form a mass-flow deposit subsequent to each explosion (LORENZ et al. 2002, LORENZ and KURSZLAUKIS 2007). In Patagonia, Argentina (Plate IV, 2) for instance in the center of the Patagonian mafic Cenozoic plateau lava fields newly discovered diatremes stand about 100m above the surrounding plane exposing the lower dia- tremes of former phreatomagmatic volcanoes and their feeding dyke systems (MARTIN et al. 2005). Diatremes them- selves are cone-shaped volcanic structures cut into pre-eruptive rocks. They are up to 2.5 km deep and up to 1-2 km in upper diameter. They are filled by clastic debris, subsided larger blocks and frequently intrusive rocks. The volume of the diatreme fill is about the same as that of the thinly bedded tephra ring and distal ash deposits. Thus, diatremes form an important part of the maar-diatreme volcano. The rather regular cone-shaped diatremes continue at depth into a root zone. This rootzone is irregular in shape and overlies the magmatic feeder dyke of the volcano (LORENZ and KURSZLAUKIS 2007). Uprising magma from the underlying feeder dyke into the diatreme root zone intrudes the clas- tic debris in the diatremes, inflates them and mingles with the debris to commonly form subterranean peperite (LORENZ et al. 2002, LORENZ and KURSZLAUKIS 2007). In “soft substrate” environment (Figure 5.9) maar volcanoes are broad and underlain by “champagne glass” shape diatremes (LORENZ 2003, SOHN and PARK 2005, AUER et al. 2007). In contrast the crater wall of maar volcanoes that erupted through “hard rock” environment (Figure 5.9) will be steep, fulfilled with volcaniclastic delta deposits and underlain by deep diatremes (LORENZ 2003, SOHN and PARK 2005, AUER et al. 2007). Maar-diatreme volcanoes are asso- ciated with any magma type involved in volcanism. The West Eifel Volcanic Field, Germany, represents a classic example of how pre-existing country rock structures influenced and controlled the position and emplacement behaviour of rising magma in the uppermost crust (SCHMINCKE 1977, LORENZ and BÜCHEL 1980, LORENZ 1984, BÜCHEL 1993). The West Eifel maar-diatreme volcanoes were formed by phreatomagmatic activity on the intersections of mostly basaltic dykes with local water bearing faults or joints especially, but not exclusively, underneath fault-controlled valley floors (Figure 5.10) (LORENZ 1984). If meteoric water was not available, scoria cones were formed on dykes by magmatic activity, without the formation of diatremes (LORENZ 1984). Kimberlites and carbonatites also seem to follow an emplacement pattern that appears to be controlled by coun- try rock structures (DOWNES et al. 2007). While their final emplacement mechanism is still under discussion (the major diatreme-forming process is thought to be related either to the expansion of a juvenile gas phase or to the inter- action of the rising magma with ground water), work published mainly over the last two decades has shown that the position of many of these economically important pipes is to at least some extent controlled by crustal discontinu- ities. Deep-seated shear zones, faults, mobile belts, transform faults, arch-style uplifts of pre-existing basement structures and the general stress field at the time of emplacement of these magma types have been quoted as being responsible, or at least as having influenced, the emplacement of pipes and dykes (JAQUES and MILLIGAN 2004, JONES and CRAVEN 2004, GRIFFIN et al. 2005). Of interest, particularly for exploration purposes, is the extent to which the 111position and emplacement of a kimberlite pipe is controlled by country rock structures. In addition, it is important for mining purposes to define the interaction of the pipe shapes with inhomogeneities in the country rock. The mor- phology of root zones of maar-diatreme volcanoes seems to be particularly susceptible to country rock faults and joint patterns. Depending on magma type involved, locality and state of erosion, maar-diatreme volcanoes may be of 112 Figure 5.9. Theoretical model of maar crater evolution in hard and soft substrate environment based on study of the Fekete- hegy maar volcanic complex in western Hungary (after AUER et al. 2007)economic relevance. The economically most relevant maar-diatreme volcanoes are diamondiferous kimberlite and lamproite pipes which occur on all major cratons. In the West Eifel, Germany, some dia- tremes underlying maar craters channel CO 2 towards the surface which may be used in fizzy water or soft drinks, or for purely indus- trial purposes. The pyroclastic rocks of diatremes and maar tephra rings may represent material suitable for use as road metal and sim- ilar purposes. Tephra ring deposits The tephra ring successions are commonly interpreted to be deposit- ed predominantly from pyroclastic density currents such as base surges (Plate IV, 3) and deposited through a gradual loss of energy as well as input of fall material into the passing base surge currents (DELLINO et al. 1990, 2004a, 2004b, DELLINO 2000, DELLINO and LA VOLPE 2000, The interbedded scoriaceous fall deposits are common in tephra ring succes- sions (Plate IV, 4) and are interpreted to be a result from an intermittent change in eruptive style, where water availability to fuel magma–water interaction was suppressed (HOUGHTON et al. 1999). In the distal sections, the base surge dominated beds are gradually replaced by reworked dm-thick beds of fine ash deposited from syn-volcanic reworking by debris flows and/or hyperconcentrated mass flows (Plate IV, 5) (LAJOIE et al. 1992, SOHN 1996, V AZQUEZ and ORT 2006). Facies variation in tuff ring sequences have been recognized characteristic changes. From the Crater Elegante in Mexico proximal sandwave bed facies replaced by massive facies in the medial regions, and the by planar bed facies in the most distal regions (Figure 5.11) (WOHLETZ and SHERIDAN 1979). A facies vari- ation of a single surge unit from the Hopi Butte recognized different trend; having disorganized beds in the rim section, stratifified beds and sandwave beds in the medial sections,and plane parallel beds in distal regions (Figure 5.12) (V AZQUEZ and ORT 2006). In the distal areas the bedding tends to be subhorizontal, with no dramatic undulations of bed contacts. The number of impact sags or accretionary lapilli-bearing beds in the distal sections of the tephra ring succession become less common. The beds are laterally continuous over tens of metres. The common inter-fingering of phreatomagmatic deposits with syn-volcanic reworked volcaniclastic sediments in the tephra ring successions generally indicates an ongo- ing remobilisation of freshly deposited tephra during the eruption. This remobilisation of freshly deposited tephra is best to be inferred to be similar to sheet ero- sion on the flank of the volcano. Mud flows and debris flows in the distal areas form tabular beds mantled by predomi- nantly base surge generated tephras. The proportion of reworked tephra in the vol- caniclastic succession is increasing with distance from the vent. 113 Figure 5.11. Facies distribution around the Crater Elegante, Mexico (after WOHLETZ and SHERIDAN 1979) Figure 5.10. Theoretical model of the formation of maar volcanoes along hydrologically active zones in the Eifel, Germany (after LORENZ 1984) Figure 5.12. Facies distribution in measured stratigraphic logs from the Hopi Butte, Arizona (after VAZQUEZ and ORT 2006: fig. 3)114 Maar lakes When the phreatomagmatic explosions of a maar-diatreme volcano finally come to an end, the crater fills up with water. The resulting maar lakes are deep in relation to their diameter and damned from the surroundings by the ejected material (crater wall)(Plate IV, 6). This special architecture effects the lake and its sediments. These lakes trap the mate- rial of an extreme small catchment area (WILDE and FRANKENHAUSER 1998, ZOLITSCHKA et al. 2000, SCHARF et al. 2001, MINGRAM et al. 2004a, 2004b, KOTTHOFF and SCHMID 2005, SABEL et al. 2005). Allochthonous clastic material reaches the crater mainly as turbitity currents originating from the crater rim (ejected material) (DROHMANN and NEGENDANK 1993). The autochthonous sediment in maar lakes is often dominated by algal material. Algal bloom layers alternate with the background sediment layers creating laminated deposits (MINGRAM 1998, BELIS et al. 1999, BRUKNER-WEIN et al. 2000, SCHABETSBERGER et al. 2004). And maar lakes often develop a meromiktic division of their water column provid- ing in this, way exceptional conditions for the preservation of sedimentary structures as well as fossils. During the post- eruptive history of a maar volcano, landslides, crater-wall collapses and the initiated volcaniclastic debris flows and tur- bidity currents form a typical lacustrine succession in the maar lake (Plate IV, 7). This destructive period is often accom- panied by a lengthy period of quiet, during which the deposition of suspended material may produce laminated sequences. These laminites are often characteristic of the palaeoenvironment, and therefore maar volcanoes are common- ly considered to be excellent sites in which continental depositional records can be well preserved (VAZQUEZ et al. 2004, CRAUSBAY et al. 2006, GARCIN et al. 2006). Large tuff rings and maars may host significant thickness of lacustrine sediments accumulate over long time (thou- sands of years) in their crater. In the western Pannonian Basin for instance at least 3 maar crater are known where few tens of metres thick lacustrine succession is preserved (MARTIN and NÉMETH 2004). Such lacustrine sediments are wet, and having significant pore space, and over time the thickness of the succession could drop significantly, that may cause significant subsidence of the crater fillings (SUHR et al. 2004). Study of such lacustrine successions bear many evidences of such processes (SUHR et al. 2004), in a form of convolute and distorted beddings, water escape structures or unusual blocks upon each other (NÉMETH et al. 2002, CSILLAG et al. 2004). After erosion and exhumation of the lacustrine suc- cession the position of the lacustrine succession maybe used to estimate the original position of the syn-eruptive land- scape on what the volcano developed. However, the gradual subsidence and compaction of the crater filling succession must be taken in account for correct estimate of the syn-volcanic surface estimates (Figure 5.13). In any case when sub- sidence and/or truncation of the lacustrine succession are recognised, the position of the syn-eruptive surface could be significantly higher than it maybe expected from the pure basis of the size and sedimentary features preserved in the suc- cession. In case a phreatomagmatic (e.g. maar) volcano forms, its crater may reach up to a few km in diametre. Such a size of crater could be filled with a significant amount of water (tens of metres range). In such crater lakes, in case of rejuvena- tion of volcanic activity, eruption may take place purely subaqueous to emergent (NÉMETH et al. 2007). After erosion vol- caniclastic sediments could accumulate in the crater lake forming complex sedimentary facies associations similar to those forming in normal subaqueous sedimentary basins (NÉMETH et al. 2007). To distinguish between sedimentary environments purely related to the syn-eruptive basin set- tings from sedimentary environments created by volcanic eruptions is important to reconstruct the syn-volcanic land- scape evolution of the region. Only the correct identifica- tion of 3D relationships between volcanic facies associated with the volcanic complexes allow correct reconstruction of the position of the syn-volcanic landscape. Geopysical anomalies over maars and diatremes vary in their character and do not provide definitive evidence for a phreatomagmatic origin (BÜCHEL 1987, SCHULZ et al. 2005, CASSIDY et al. 2007). Many known maars and diatremes are buried by post-genetic sediments. Geophysical methods resulted in their initial discovery and subsequent drilling provided geologic samples, which confirmed their phreatomagmatic origin. Interpretation of a single geophys- ical data set over a suspected maar or diatreme structure can be ambiguous. When combined, however, with complemen- tary geophysical methods and the existing database over other known maar or diatreme structures, a more definite Figure 5.13. Posteruptive subsidence and deformation of Miocene lignite bearing sediments above the Oligocene Kleinsaubernitz maar-diatreme volcano, eastern Saxony, Germany (after LORENZ 2007: fig. 4)115 assessment can be made. The most notable geophysical signature associated with maars or diatremes is a negative grav- ity anomaly (SCHULZ et al. 2005). These gravity lows are generally circular and cover the whole structures. They are due to lithological and physical changes associated with the preatomagmatic explosion. In well-preserved maar structures, low-density sedimentary infill of the topographic depression of the crater contributes to the gravity low. In general, mag- netic anomalies associated with maar or diatreme structures are more complex than gravity anomalies (SCHULZ et al. 2005). Their reasons are very complex intrusion processes in the diatreme or in the sedimentary infilling of the maars. Also the development of spatter cones into the maar crater can cause a magnetic anomaly.The presence of fluids in explo- sion-induced fractures and pore spaces of the maar and diatreme rocks leads to decreased resistivity levels that can be mapped effectively by various electrical methods (BRUNNER et al. 1999). Reflection seismic surveys allow for detailed imaging of maar structure morphology (Figure 5.14). Well logging methods are very useful for the detailed investigation of drill holes in maar and diatreme structures. Peperite associated with small volume phreatomagmatic volcanoes Peperite results from interaction between magma and wet sediment and exhibits a range of complex textures (WHITE et al. 2000, SKILLING et al. 2002, WHITE and HOUGHTON 2006). Peperite is common in many geological settings where magma comes in contact with wet sediment (SKILLING et al. 2002). The presence and recognition of peperite indicate contemporaneous volcanism and sedimentation, and they provide insight into the nature of subsurface magma transport and host-sediment properties at the time of eruption (SKILLING et al. 2002). Peperite is common along the margin and adjacent to intrusive bodies and where lava flows travelled through wet sediments (SKILLING et al. 2002). Recognition of peperite-forming processes in a sub-volcanic region of a phreatomagmatic volcano is given from several sites (HOOTEN and ORT 2002, MCCLINTOCK and WHITE 2002). Peperites commonly form when magma intrudes wet unconsolidated sediment. The host can be any clastic sediment, and interactive phenomena occur in great variety under a wide range of physical conditions. In general, peperite are con- sidered to be very common in arc-related and other volcano-sedimentary sequences associated with composite volcanic systems. In such settings peperite and associated volcaniclastic sediments may form large volume of deposits of a very diverse textural and compositional range from rhyolite to basic rock types. Peperites and their associated volcaniclastic environment are important in reconstructing palaeoenvironments. It has been recently recognized that peperite and their associated pyroclastic environment are as common in small volume phreatomagmatic vent-filling deposits and/or along the contacts between sediment and mafic, generally small-volume intrusions and lavas as in more complex stratovolcanes and or caldera volcanic systems (MARTIN and NÉMETH 2007, NÉMETH and MARTIN 2007). In the Pannonian Basin, a wide range of peperite has been described lately in phreatomagmatic, small-volume mafic volcanic systems (Plate V, 1). The recognition of the existence of peperite in small-volume, phreatomagmatic volcanoes and their associated intrusive/effu- sive products highlights the importance of identification of such textures and the interpretation of their meaning in these types of volcanoes (MARTIN and NÉMETH 2007). Peperites are more common in association with small-volume terrestri- al phreatomagmatic volcanoes than it has been considered previously, especially in volcanic fields, which developed in Figure 5.14. Interpretative reflection seismic profiles of two maar from Germany; a) Baruth, b) Messel (after SCHULZ et al. 2005: fig. 5)a region where surface and sub-surface aquifers provided substantial water and/or water saturated sediments to sustain phreatomagmatism. Peperite is often described on the basis of the juvenile clast morphology as blocky or fluidal, but there is a big range in shapes that have to be taken in to account as well (BUSBY-SPERA and WHITE 1987). The shape dif- ferences are often explained by the simple textural differences of the host deposit as fine and/or coarse grained texture that may produce globular (Plate V, 2) and/or blocky peperite respectively (BUSBY-SPERA and WHITE 1987). However, examples from the phreatomagmatic volcanic fields in the Pannonian Basin show that this simple classification is not the only controlling parameter (MARTIN and NÉMETH 2007). Peperite in association with phreatomagmatic volcanoes can develop when feeder dykes intruded water-saturated sediment infilling 1) maar or tuff ring basins and/or 2) vent zones and/or 3) lava flow outpouring from such volcanoes (Figure 5.15). However, these three types of peperite may suggest strikingly different environmental reconstruction and therefore their interpretation may imply different geomorphologi- cal reconstruction. In the western Pannonian Basin (Hungary) lava capped buttes, e.g. Kissomlyó (MARTIN and NÉMETH 2005) are commonly accompanied with a great variety of peperite that indicate wet environment into which the lava flow erupted. However, the reconstruction of the position of the lava flows need to be established before any geomorpholog- ical conclusion. In the western Pannonian Basin, most of the lava flows erupted inside of a crater of a phreatomagmatic volcano, and therefore the presence of peperite only means that the crater zone and not necessarily the surrounding syn- volcanic landscape was “wet” and water saturated. The identification of peperite in such settings therefore does not guar- antee the reconstruction of the eruptive environment nor the establishment of the syn-volcanic landscape position onto the volcano erupted. “Dry”magmatic processes associated with phreatomagmatic volcanism Maar-diatreme volcanoes are created by phreatomagmatic eruptions arising from a contact of ascending magma with ground water in available aquifers. If this contact is eliminated magma continues its ascent towards the surface giving rise to a wide range of volcanic forms and products associated closely with maar-diatreme volcanoes. Usually this hap- pens in the advanced or closing stage in evolution of the volcano. However, in arid regions the maar-forming eruptions were often preceded by effusive and/or Strombolian activity (Plate V, 3) (GUTMANN 2002). Various scenarios are possi- ble and factors controlling changes in the eruption style are understood poorly. Phreatomagmatism might be inhibited unless the magma flux is low relative to the rate of water supply and unless the top of the magma column has subsided, probably below the water table. A general evolutionary sequence from hydromagmatic eruptions during formation of the maars, through Strombolian eruptions of the post-maar scoriae and ashes, and finally to the post-maar lavas (Plate V, 4) appears to reflect the declining influence of magma – ground water interactions with time, which process widely docu- 116 Figure 5.15. Schematic representation of peperite types identified from the Neogene phreatomagmatic volcanic fields of Hungary (after MARTIN and NÉMETH 2007: fig. 13)mented in many volcanic fields such as Hopi Buttes (Arizona) (WHITE 1991b). The general time sequence of eruption events (from wet to dry explosive processes) and the distribution of volcanic vent type (dry and wet) in regard to the geo- morphology of the pre-volcanic landscape is well established, having phreatomagmatic volcanoes in low lying areas and scoria cones in highlands (WHITE 1991a). Transient hydromagmatic events e.g. occurred relatively late in the 1975 erup- tion sequence of the Tolbachik volcano in Kamchatka (DOUBIK and HILL 1999). Maar rims are thin and composed of eas- ily erodable tephra. Intermittent phreatomagmatic acitivity due to sudden ground water access to the volcanic conduit in the vaning stage (low magma supply rate) of the eruption have also been described in large ocean islands such as Hawaii Kilauea (DZURISIN et al. 1995). The most widespread of these deposits at Kilauea (Uwekahuna Ash Member) is a few metres thick basaltic surge and fall deposit (DZURISIN et al. 1995). These few metres thick deposits inferred to be a result of two major pyroclastic surges, each preceded by unusually vigorous lava fountaining from a vent near the volcano’s summit (DZURISIN et al. 1995). Simultaneous magmatic and phreatomagmatic explosive events in the same volcano have been recognized from tuff ring sequences of many volcanoes. For instance, a series of alternating phreatomagmatic (“wet”) and magmatic (“dry”) basaltic pyroclastic deposits forming tuff rings, such as the Crater Hill tuff ring in New Zealand, contains unit(s) which can only be interpreted as the products of mixing of ejecta from simultaneous wet and dry explosions at different por- tions of a multiple vent system (HOUGHTON et al. 1999). In other type of mixed deposit has been recognized from the Eifel, where documented mixed (wet and dry) basaltic pyroclastic deposits interpreted to represent mixing from two point sources (e.g. vents) of quite different but stable character (HOUGHTON and SCHMINCKE 1986, 1989). These two well-doc- umented site highlights the complexity of eruption dynamics of a predominantly “wet”, mafic volcano. Magma, that has not reached the surface, appears as dykes, sills and/or plugs in diatreme/maar filling. Some of them may represent feeders to surficial activity. Outpourings of lava feed up lava flows and/or lava lakes in maar depressions, often in fully subaqueous environment. In such case pillow lavas, hyaloclastite breccias and/or peperite breccias are pres- ent. Phreatomagmatic eruptions of the Surtseyan-type due to interaction of ascending magma with water in the maar lake give rise to palagonite tuff cones. Eruption rate and water depth are factors controlling Surtseyan-type eruptions and tran- sition towards Strombolian-type eruptions. Lava flows and lava lakes provide an excellent opportunity to study evolution of jointing and its relationship to the form of the lava body. They are also good objects to calibrate and compare Quaternary dating methods, remote sensing methods, and rates of geomorphic processes. Surtseyan volcanism Surtseyan eruptions (Figure 5.16) are characterized by interaction of a fluid erupting magma with abundant external water (KOKELAAR 1983, WHITE and HOUGHTON 2000). Commonly they start in shallow subaqueous environments (Figure 5.17) where accumulating tephra forms a mound-shaped volcano prior to emergence (WHITE 1996, WHITE and HOUGHTON 2000). During this subaqueous stage density currents play a significant role in shaping the edifice and imparting charac- teristic bedding features (WHITE 1996, SMELLIE and HOLE 1997, MUELLER et al. 2000, WHITE 2000, MARTIN 2002). Emergent to subaerial Surtseyan eruptions generally produce cone-shaped edifices by accumulation of wet fall deposits with subordinate density currents (VERWOERD and CHEVALLIER 1987, SOHN and CHOUGH 1993, SOHN 1995, COLE et al. 2001). Surtseyan deposits typically consist almost entirely of glassy fragments formed by fragmentation of the erupting magma, and lack a significant country- rock component (KOKELAAR 1983, WHITE and HOUGHTON 2000). This is an impor- tant difference from deposits of maar- forming eruptions, and it indicates that in Surtseyan eruptions fragmentation occurs at very shallow levels in the edifice or/and as the magma emerges from it. If or when the erupting magma no longer encounters water (e.g. by enclosure of an emergent vent, or isolation from/depletion of ground water), both Surtseyan and maar- forming eruptions may transform to Strombolian or Hawaiian ones. The sub- aqueous phases of the eruption of an emergent volcano, such as the well-docu- mented Pahvant Butte eruption initiated few tens of metres (85 m in the case of 117 Figure 5.16. Eruption model of a Surtseyan eruption (after KOKELAAR 1983)118 Pahvant Butte) beneath the water surface produced a broad mound of tephra (Figure 5.18)(WHITE 1996). A variety of distinctive lithofacies identified from Pahvant Butte allows reconstruction of the eruptive and depositional processes active prior to emergence of the volcano above lake level (WHITE 1996). Early in the eruption subaqueous tephra jet- ting from phreatomagmatic explosions discontinuously fed inhomogeneous, unsteady, dilute density currents (WHITE 1996). Dunes and crossbeds which are better developed upward in the section of Pahvant Butte resulted from interaction between sediment gravity flows and sur- face waves triggered as the explosion-generated pressure waves and eruption jets impinged upon and occasionally breached the surface (WHITE 1996). Intermingling of (a) tephra emplaced after brief transport by tephra jets within a gaseous milieu and (b) laterally flowing tephra along vent margins during parts of the eruption in which episodes of continuous uprush produced localized water- exclusion zones above a vent. Mass flow deposits formed by disruption and remobilization of mound tephra. Intermittent, explosive magma– water interactions occurred from the outset of the Pahvant eruption, with condensation, entrainment of water and lateral flow mark- ing the transformation from eruptive to “sedimentary” processes leading to deposition of the mound lithofacies (WHITE 1996). Surtseyan-style eruptions have been observed in marine environment (THORARINSSON et al. 1964, 1967) and crater lakes (NÉMETH et al. 2006). The observations confirmed that two type of explosive activity accompanied with the eruption (KOKELAAR 1983, WHITE and HOUGHTON 2000); 1) repeated continuous particle up- rush, repeated in about up to 20 minutes intervals, and 2) explosive bursts forming dark coloured, ash charged cock’s tail jets repeated in few minutes intervals. Surtseyan-style volcanoes are commonly develop in the rift edges of ocean islands. Along the rift axis scoria cones form, and in the low-lands, near the sea level, phreatomagmatic volcanoes as well as Surtseyan-style volcanoes form such as the 1957-58 eruption of Capelinhos in Azores (Figure 5.19) (MACHADO et al. 1962, WATERS and FISHER 1971, COLE et al. 2001), or the 1913 eruption sites on Ambrym, Vanuatu (NÉMETH and CRONIN 2006). Surtseyan-style eruptions are also common in island arc settings such as the Izu-Bonin arc in Japan. In such settings, volcanic islands (tuff cones) can form Figure 5.18. Cross section (a) of Pahvant Butte with identified lithofacies (b) (M1-4). The structure and morphology of the butte probably adoptable to other subaqueous phreatomag- matic volcanoes' architecture (after WHITE 1996) Figure 5.17. Eruption model of the Pahvant Butte. A) initial subaqueous phase, B) subaqueous fall formation, C) emergent phase,eruption cloud breaches the lake surface (after WHITE 1996)119 quickly but they commonly completely vanish by wave action, such as the Graham Island (Ferdinandia) in the Channel of Sicily in 1831 (PlateV, 5). Surtseyan eruption also can take place in caldera lakes such as the 2005 Ambae (Plate V, 6) eruption in Vanuatu (NÉMETH et al. 2006), where the 50 m high tephra cone erodes quickly after the eruption ceased. Volcanic hazards of monogenetic volcanoes Hazards associated with maar eruptions are: volcanic earthquakes (up to c. M: 4-5), possibly several 1000 individ- ual eruptions, eruption clouds rising to maximum heights of economic air travel, ejection velocities of tephra clasts of up to 400 m/s, ejection distances of ballistic clasts up to 4 km; size of ejected clasts up to 8 m, base surges travelling up to several km and with time building a tephra ring of a height up to 100 m and of a radius of up to 4 km (measured from centre of crater), thin distal tephra falls extending to more than 100 km, syn- and post-eruptive slumps and lahars inside and in part also outside the crater, destruction of buildings and transport lines within a radius of up to 5-6 km (LORENZ 2007). Associated formation of the maar crater floor and underlying diatreme results in subsidence of coun- try rocks, tephra, and buildings to depths of possibly 1000-2000 m (SUHR et al. 2004). In addition, recent studies have shown that there are hazards associated to recurrence of activity within volcanic fields but also in single maars. Volcanic hazard studies of a volcanic field commonly target to understand the recurrence rate, the eruption frequency and style of eruptions may occur in the future (HO 1992, CONNOR and HILL 1995, HO and SMITH 1998, CONNOR et al. 2000, CRONIN et al. 2001, CRONIN and NEALL 2001, EDBROOKE et al. 2003, HURST and SMITH 2004, MAGILL and BLONG 2005a, 2005b, MAGILL et al. 2006). Figure 5.19. Map of the Capelinhos volcano developed in the edge of a volcanic island, Azores (after COLE et al 2001)120 References ARANDA-GOMEZ, J. J. and LUHR, J. F. 1996: Origin of the Joya Honda maar, San Luis Potosi, Mexico. — Journal of Volcanology and Geothermal Research 74 (1–2), pp. 1–18. ARANDA-GOMEZ, J. J., LUHR, J. F. and PIER, J. G. 1992: The La-Brena-El-Jaguey-Maar Complex, Durango, Mexico.1. 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W. 2000: Eifel maar lake sediments: Calendar-year archives of past environments. — Terra Nostra 6, pp. 545–549. 125126 Chapter 5 Location map 1 — Hopi Butte, Arizona, USA 2 — Lamongan, Java, Indonesia 3 — Pannonian Basin, Central Europe 4 — Snake River, Idaho, USA 5 — Ukinrek maar Alaska, USA 6 — Rininahue, Chile 7 — Deccan, India 8 — Etna, Italy 9 — Iceland 10 — Eastern Australia 11 — Hawaii 12 — Tihany, Hungary 13 — Füzes-tó, Hungary 14 — BBHVF, Hungary 15 — La Brena maar, Mexico 16 — Paricutin, Mexico 17 — Tarawera, New Zealand 18 — Patagonia, Argentina 19 — Pali Aike, Argentina 20 — Chubut, Argentina 21 — Payunia, Argentina 22 — Ceboruco, Mexico 23 — West Eifel, Germany 24 — Jeju Island, Korea 25 — Pula maar, Hungary 26 — Messel maar, Germany 27 — Baruth maar, Germany 28 — Ság-hegy, Hungary 29 — Hajagos, Hungary 30 — Kissomlyó, Hungary 31 — Crater Elegante, Mexico 32 — Tolbachik, Russia 33 — Crater Hill, New Zealand 34 — Surtsey, Iceland 35 — Pahvant Butte, Utah, USA 36 — Capelinhos, Azores, Portugal 37 — Ambrym Island, Vanuatu 38 — Graham Island, taly 39 — Ambae Island, Vanuatu127 Chapter 5 Plate I 1. MrSID satellite image of the Hopi Butte volcanic field where about 300 monogenetic volcanic vents have been identified. The map view is about 80 km wide. 2. Typical scoria cone from Payunia, Mendoza, Argentina. 3. A well developed tuff cone in the western edge of Ambrym, Vanuatu. 4. Low tuff ring from Payunia, Mendoza, Argentina. 5. A maar with lake from the Pali Aike volcanic field, Patagonia, Argentina. 6. Typical tephra ring sequence consists of alternating base surge and phreatomagmatic fall beds from a maar rim in the Pali Aike Volcanic Field, Argentina. 7. Deep and wide crater of the Crater Elegante maar in Sonora, Mexico.128 Chapter 5 Plate II 1. Exposed diatreme from Chubut, Argentina. 2. Scoria cone and lava field in the maar crater of the La Brena maar, in Durango, Mexico. 3. Large scoria cone of Paricutin in Michoacan, Mexico. 4. Intracontinental lava shield volcano from Al Haruj, Libya. 5. Lava spatter-dominated scoria cone chains in the Pali Aike Volcanic Field, Argentina. 6. Large tumuli from the Al Haruj, Libya.129 Chapter 5 Plate III 1. Lava spatter deposits preserved the steep flank of a volcanic cone of the Al Haruj Volcanic Field, Libya. 2. Collapsed spatter cone fragments carried away on a lava flow as documented from a cone complex of the Crater Basalt Volcanic Field, Chubut, Argentina. 3. Typical scoria cone deposit of a scoria cone from scoria cone nearby Ceboruco, Western Transmexican Volcanic Belt, Mexico. 4. Welded agglutinate succession from Al Haruj, Libya. 5. Completely welded and exhumed interior of a scoria cone from Al Haruj, Libya. 6. Chain of scoria cones near Ceboruco, Nayarit, Mexico.Chapter 5 Plate IV 1. Maar lake in Rininahue, Chile. 2. Exposed diatrema from southern Slovakia, the sharp plain is the contact between diatreme (left) and host (right) rocks. 3. Typical base surge beds from the 1913 eruption of west Ambrym, Vanuatu. 4. Interbedded scoriaceous fall deposit in the 1913 phreatomagmatic succession of west Ambrym, Vanuatu. 5. Reworked volcaniclastic succession in distal areas of a tephra ring of the 1913 eruption site of west Ambrym, Vanuatu. 6. Large maar lake (Potrok Aike) from the Pali Aike Volcanic Field, Patagonia, Argentina. The tephra ring is almost entirely eroded. 7. Soft sediment deformation in maar lake sediment of the Pula maar, western Hungary.131 Chapter 5 Plate V 1. Blocky peperite from the Hajagos maar in western Hungary. 2. Globular peperite along a sill from the Ság-hegy phreatomag- matic volcano. 3. Scoria cone half section exposed in the crater wall of the Crater Elegante maar in Sonora, Mexico. 4. Transition from phreatomagmatic succession (base of hill along the cliff of the Snake River valley) to strombolian scoria cone units (capping topmost part of hill) from the Sinker Butte, Western Snake River Plain, Idaho. 5. Two paintings of the eruption of the Graham Island in 1831 in the Channel of Sicily, Mediterranean Sea; on (A) white eruption cloud probably caused by steam is visible with dark jets closely resembling cocks’ tail jets reported from Surtseyan eruptions. On (B), lava fountaining is visible, possible representing the late stage of the eruption of the emergent volcano, where no or just limited magma-water interaction took place. 6. Surtseyan eruption in Lake Voui in Ambae Island in December, 2005. Polygenetic volcanism Polygenetic volcanism (Figure 6.1) is considered to be volcanism that is associated with long lasting and usually com- plex eruption history of a volcanic system (DA VIDSON and DE SILVA 2000; LIPMAN 2000; WALKER 2000). Such volcanic systems can either consist of a single volcanic edifice, or where slight shifting of position of the active pathway of magma to the surface occurs a nested and multiple edifice complex can result. In general, we view a volcanic system as polyge- netic if successive magma batches causes eruptions in more or less in the same place (DA VIDSON and DE SILVA 2000). In this way, the term can be slightly misleading because many relatively small volume eruptions may recur in the same place and may result complex, nested volcanic systems such as those found in many long lived basaltic volcanic fields (HOUGHTON and SCHMINCKE 1989; HOUGHTON et al. 1996; CONNOR and CONWAY 2000; VESPERMANN and SCHMINCKE 2000). Many of relatively small volume volcanoes in a basaltic volcanic field can exhibit very complex volcanic archi- tectures, and in many cases their eruption history shows signs of multiple and recurring activity over long periods of time (AUER et al. 2006). However, it is rare for small volume basaltic volcanoes to stay active for longer then few months (SELF et al. 1980; LUHR and SIMKIN 1993), in contrast to true polygenetic volcanoes commonly forming composite volcanic edi- fices that are active over thousands of years. The transition between small volume but complex basaltic volcanoes and long-lived but relatively small volume composite volcanoes is rather gradual and it has been described from many vol- canoes from Central America (ABRAMS and SIEBE 1994; MCKNIGHT and WILLIAMS 1997). Polygenetic volcanoes can occur in many different edifice forms as well as being very different in duration of activ- ity. In this way a large, long-lived shield volcano associated with hot spot magmatism is also classified as polygenetic (Plate I, 1), and can result in accumulation of large volumes lava fields, with very diverse geochemical signatures, reflecting slight chemical changes in the rising melt over long periods of time. A good example of these types of vol- canoes is the islands of Hawaii where hot spot activity formed four coelescing broad shield volcanoes over the last mil- lion years of activity (FREY et al. 1991; MOORE and CLAGUE 1992; MOORE 1992). The activity resulted in the accumulation of diverse composition and texture pahoe- hoe lava fields (KESZTHELYI and SELF 1998; CROWN and BALOGA 1999; BYRNES and CROWN 2001). Such shield volcanoes can also have a complex architecture and the volcanic edifice can contain a reasonable volume of pyro- clastic material resulting from occasional explosive activ- ity (MCPHIE et al. 1990; DZURISIN et al. 1995; MASTIN and WITTER 2000). Their reworked counterpart may also form extensive volcaniclastic aprons around the main lava shield. Due to gradual growth of a shield volcano, and their relatively unsupported seaward edifice constructs (MARTI et al. 1997; MASSON et al. 2002; MORGAN and CLAGUE 2003), such volcanoes may also develop exten- sive volcaniclastic aprons in the sea floor, commonly as a result of occasional volcano collapse (ANCOCHEA et al. 1994; GARCIA and HULL 1994; MOORE et al. 1994; W ATTS 133 Chapter 6 Figure 6.1. Variation of volcanic landforms as a function of their size and morphological parameters (from DA VIDSON and DE SILVA 2000: p. 664, fig.1)and MASSON 1995; MCMURTRY et al. 1999; CERVELLI et al. 2002; ENGELS et al. 2003; COOMBS et al. 2004). Large shield volcanoes can also form large volume of hyaloclastite due to their repeated effusive eruptions in the sea floor (Plate I, 2) (QUANE et al. 2000; WALTON and SCHIFFMAN 2003; COOMBS et al. 2004; SCHIFFMAN et al. 2006) and littoral cones (JURADOCHICHAY et al. 1996) where lava and sea water contact take place. Such hyaloclastite aprons can be drifted away by currents and resedimented around the volcanic edifice (FURNES and STURT 1976; RUIZ et al. 2000). Since a lava shield volcano can be active over hundreds of thousands of years, such sedimentary units can form a significant volume of sediments (GEE et al. 2001; KRASTEL and SCHMINCKE 2002; MITCHELL et al. 2002). Such sediments can be preserved in a sedimentary record. In ancient sedimentary sequences thick piles of resedimented hyaloclastite marks long lasting and polygenetic evolution of a lava shield. Lava shield volcanoes are also common in subaerial settings in intracontinental tectonic settings, for example (Plate I, 3) in northern Africa (ADE-HALL et al. 1974; FRANZ et al. 1997, 1999; WILSON and BIANCHINI 1999; KENEA et al. 2001; NÉMETH 2004; MARTIN and NÉMETH 2006), along the Snake River Plain (Plate I, 4) in Idaho (GREELEY 1982; GEIST and RICHARDS 1993; CUMMINGS et al. 2000; GODCHAUX and BONNICHSEN 2002) or in Eastern Australia (FREY et al. 1978; SUTHERLAND 1983; JOHNSON 1989; OREILLY and ZHANG 1995). Volcanoes in these environments can also be long lived and usually their position marks major magma rise paths in an intracontinental vol- canic field. The large accumulation of lava flows in such volcanic fields (STEPHENSON and GRIFFIN 1976; STEPHENSON et al. 1996) may also have associated explosive phases of activity (Hawaiian- or Strombolian-style eruptions) as well as occasional phreatomagmatic explosive activity (NÉMETH et al. 2003; PEREGI 2003; NÉMETH 2004). Lava shields in such settings adjacent to a lacustrine basin may give rise to lava deltas and hyaloclastite successions (Plate I, 5) may devel- op in the areas where lava flows enter fluvio-lacustrine networks (GODCHAUX and BONNICHSEN 2002; HUGHES et al. 2002; WOOD and CLEMENS 2004). Similar facies may develop where lava enters shallow marine environment (BEHNCKE 2004). Although intracontinental lava shield volcanoes can be volumetrically the same as those forming ocean islands, their size is commonly a magnitude smaller. Intracontinental flood basalt volcanism is also inferred to be long lived (tens of thousands of years of activity), and commonly forms large volumes of phreatomagmatic tephra as described from the Jurassic Mawson Formation in Antarctica (ROSS and WHITE 2005; MCCLINTOCK and WHITE 2006) or Karoo from South Africa (ROSS et al. 2005). Such phreatomagmatic pyroclastic rocks may form as a result of long lasting phreatomagmatic eruptions that may generate so called phreato-cauldrons, which leave large volume of deposits (WHITE and MCCLINTOCK 2001). Composite or strato-volcanoes are a major type of polygenetic volcanoes. They occur predominantly in subduction zones, however, intraplate volcanoes of similar architecture are also common such as those forming a Cenozoic vol- canic chain in Eastern Australia. Composite volcanoes are those that erupt over long period of time (thousands of years), and producing extensive lava flows alternating with major explosive eruptive products as well as their reworked, redeposited and resedimented counterparts (Figure 6.1). Volcanoes that form amalgamation of relatively small volume eruptive vents, commonly termed compound volcanoes, are typical of basaltic volcanic systems. Composite volcanoes are surrounded by a ring plain, a place where distal primary volcaniclastic sediments as well as lava flows and associ- ated reworked material can accumulate and form a basin fill succession, commonly intercalated with normal siliclastic sediments (PALMER and NEALL 1991). Composite volcanoes usually have a central vent (Plate II, 1) surrounded by satel- lite vents that are individual, short lived, monogenetic volcanoes. However, their feeding systems are commonly cou- pled with the central vent of the composite volcano (Plate II, 2). Such satellite vents are located in areas that are favoured by local stress fields which allows melt to reach the surface. Commonly they can change their position as the composite volcanic edifice evolves. Composite volcanoes that are active over thousands of years, can reach an equilib- rium edifice architecture, that is, a shape that is controlled by equilibrium between the total volume of the erupted prod- ucts and the erosion that acts on it especially in inter-activity periods (DA VIDSON and DE SILVA 2000). Shape of the com- posite volcano is therefore strongly controlled not only by the style of eruptions, and the texture and physical proper- ties of the eruptive products, but also by external forces such as rainfall, rainfall distribution, principal wind direction, and characteristics of the substrate that may affect the stability of the growing volcanic cone (DA VIDSON and DE SILVA 2000). In a very simplistic way, the type of volcanoes from morphological point of view can be characterised by their edifice shape (constructional versus excavational) and whether they are monogenetic or polygenetic (Figure 6.1) (DA VIDSON and DE SILVA 2000). Polygenetic volcanoes in this classification can be constructional, such as shield volca- noes and composite cones, or excavational such as calderas (DA VIDSON and DE SILVA 2000). Excavational polygenetic volcanoes, can be extremely complex, and their original landform can be truncated by uplifted resurgent part of the caldera floor (ACOCELLA and FUNICIELLO 1999; LIPMAN 2000). Such resurgence may occur more than once and, espe- cially in ancient settings, can produce a constructional volcanic landform, where geomorphic inversion is significant (ACOCELLA et al. 2000; LINDSAY et al. 2001a). Composite volcanoes are the most common type in convergent plate margins, where chains of large volume and long lived arc volcanoes can develop, and affect the sedimentation in their surroundings. Spacing of these volcanoes shows a regular pattern (STERN 2004). However, exceptions may occur, such as the Rininahue volcanic field which developed in the space between two major composite volcanoes (STERN 2004). Such exceptions may be caused by plate tectonic and 134regional stress field distribution anomalies. These types of volcanoes are usually andesitic or dacitic in composition. However, rhyolitic eruptions also occur. In the divergent plate margins composite volcanoes are rare, however, Hekla and Askja in Iceland are two large volume composite volcanoes in such settings (BALDRIDG et al. 1973; SPARKS et al. 1981; BROWN et al. 1991; GUDMUNDSSON et al. 1992; STURKELL et al. 2006). In the rifting stage of continental break up, large intraplate composite volcanoes can develop such as those associated with the East African Rift system (BOSWORTH 1987; ROCHE et al. 2001). Calderas can form during the evolution of normal composite volcanic cones. Large volumes of volatile rich rhyolitic magma can form through fractionation, that may triggers a major explosive eruption that reduce the former composite volcanic edifice to a caldera such an example is the Crater Lake in Oregon, formed by the Mazama eruption (BACON 1983; KLUG et al. 2002). Larger caldera forming eruptions are directly related to long-lived magma infiltration and shallow level magma evo- lution that produced very explosive eruptions, resulting in the formation of calderas tens of kilometres in diameter. Such caldera eruptions commonly have evolutionary stages, and the formation of calderas may be associated with intensive post caldera volcanism, that can form individual, small volume constructional volcanic edifices (LIPMAN 2000). Mega- calderas that formed few tens to thousands of km 3 volume of pyroclastic material may have erupted over long period of time, and their eruption may have taken long time are common in the Andean volcanic arc (LINDSAY et al. 2001a). Those volcanoes have only recently been reported due to their deposits and hundreds of kilometres across extension broad, flat morphology (LINDSAY et al. 2001a). Classical cone-shape composite volcanoes have certain morphological characteristics especially dacitic composi- tions. Lava domes commonly cap the summit of such volcanoes (Plate II, 3), and these may grow over decades. Gravitational instability, partially due to volatile pressure, can trigger dome collapse-induced pyroclastic flows (SATO et al. 1992). Composite volcanoes are commonly truncated by major edifice failures in the form of volcanic debris ava- lanches (Plate II, 4). In the morphological development of a composite volcano the interplay between aggradation and degradation is important (HATHWAY and KELLEY 2000). When external forces are strong (tropical climate, heavy rainfall) and the eruption frequency is relatively low, the volcanic edifice can be deeply incised, steep, and, in case of volatile rich magma involvement during eruptions, can commonly fail (Plate II, 5). In other end-member, when the degradation is slow, eruptive products can accumulate quickly and lead to gravitational failure of the volcanic edifice. The characteris- tic conical shape of active composite volcanoes such as Mayon in Philippines, or Taranaki in New Zealand (Plate III, 1), indicates an equilibrium between aggradation and degradation (DA VIDSON and DE SILVA 2000). The total volume and size of the composite volcanoes is therefore controlled by factors such as magma supply rate, magma composition, climatic forces, and perhaps the physical properties of the crust over which the volcanoes develop. Where a relatively steady and long lived magma supply exist, such as the case of many intraplate volcanoes, the edifice can grow significantly, form- ing composite volcanoes such as Ararat in Turkey amongst the largest on Earth (PEARCE et al. 1990; YILMAZ et al. 1998; ADIYAMAN et al. 2003). This given tectonic setting composite volcanoes are commonly of the same size, reflecting the common nature of the melt and the tectonic regime in which they developed. In the aggradational stage of the volcano, the total eruption product volume and eruption style (effusive versus explosive) is important in the development of a certain cone shape. Magma ascent is powered by the “hydrostatic head effect”, which is an overpressure in the magma reservoir, in turn controlled by the chem- istry of the melt (volatile content) and the tectonic stress in the crust (DA VIDSON and DE SILVA 2000; MURRAY and STEVENS 2000; RUTHERFORD and GARDNER 2000; PINKERTON et al. 2002). Final magma extrusion initiates from relatively shallow magma chambers beneath composite volcanoes (Figure 6.2) The maxi- mum height the magma may be able to reach in an “open” conduit is therefore strongly controlled by magmatic overpressure. Higher initial magmatic overpres- sure drive high eruption rates, and significant volumes of lava can erupt early in the evolution of the composite volcano, forming lava dominated, high density (“heavy”) volcanic edifice. Such an edifice can increase the lithostatic load and therefore suppresses the ability of magma to reach higher zones of the growing edifice, and, as final stage, magma cannot reach the summit vent (IDA 1999; MURRAY and STEVENS 2000). During magma ascent, storage of magma in magma chambers can allow differentiation to take place (HANSTEEN et al. 1998), and over time low density, but higher viscosity magmas can evolve. As a result the magmatic overpressure can be main- tained due to the lower density and higher volatile content of the evolved melt, even under the increasing lithostatic load of the growing edifice, but the eruptions gradually could switch from dominantly effusive to dominantly explosive. Such a process can lead to volcanic eruptions that may destabilize the edifice and cause volcano collapse. This in turn can cause decrease in lithostatic load, and returns the volcanic system to near-initial conditions. Such pulsating growth and destruction of composite volcanoes are common in the geologic record, and probably one of the factors most responsi- 135 Figure 6.2. Shematic diagram represent- ing the relation between lithostatic pres- sure (P) and eruption driving overpressure (Pex) in the magmatic system of an active volcano. (after DA VIDSON and DE SILVA, 2000: p. 671, fig. 5)ble in the development of the volcanic landforms of composite volcanoes (ANCOCHEA et al. 1990; PALMER and NEALL 1991; BEGET and KIENLE 1992; CACHO et al. 1994). Observation suggests that there are physical limits to edifice heights at around 3000 m above the base of the volcano, suggesting non-evolved compositions for the initial eruptive products. This dynamic relationship between effusion rate and style of eruption are the main controlling parameters of the shape and morphology of volcanic edifice. Equilibrium stage volcanic edifices are those that have reached their maximum height and the further morphological evolution of the volcano can only occur as aresult of major edifice failure that could reset the physical conditions. This way, this is the point when degradational processes come important. In this stage degradation can form deposits that accumulate in a ring plain that may form a stabilizing flank around the volcano to prevent major edifice failure. Edifice failure perhaps can take place in the near summit vent area that can be masked quickly by fresh tephra that has been recognized in many composite volcanoes (Plate III, 2). The shape of the volcanic edifice is largely controlled by slight shifting of the vent location. Volcanic edifices that not have a stable conduit wall can develop closely spaced vent location that can even be active simultaneously. However, the individual active vent activities can be strongly coupled. Volcanic products of composite volcanoes are very diverse. They can grow relatively quietly, through Strombolian and Hawaiian-style eruptions (GARDEWEG et al. 1998) and associated lava effusions through their central vents. Such eruptions could change to Plinian-style eruptions ower a longer time frequency and produce pyroclastic flows, common- ly controlled by major eruption column collapses. Many arc volcanoes however, grow lava domes that could collapse due to gravitational instability and/or minor explosive disruption, both leading to form block-and-ash flows such as the Unzen eruption in 1990–1995 in Japan (NAKADA et al. 1999). In inter-eruptive phases reworking and non-eruption triggered lahar and/or volcanic debris avalanche formation could take place. Since pyroclastic flow development is one of the most com- mon and dangerous phenomena associated with composite volcanoes, their origin and depositional features will be con- sidered further. Pyroclastic flow genesis Pyroclastic flows are generated byvolcanic eruptions and are considered to be very mobile, hot, and have high parti- cle concentration, move horizontally moving and be dominanted by gas-particle dispersion (FREUNDT et al. 2000; WILSON and HOUGHTON 2000). The particles in the pyroclastic current are pyroclasts, since they are generated by explosive frag- mentation of the magma and the conduit wall and then erupted through a volcanic vent. The particle-support mechanism of the pyroclastic flow is dominated by fluidisation, buoyancy, grain to grain collision, and hindered settling (SPARKS et al. 1978; WILSON 1980; CAREY 1991; FREUNDT et al. 2000; WILSON and HOUGHTON 2000). One of the most common ways in which pyroclastic flows form is by collapse of a vertical eruption column (Figure 6.3) (CAS and WRIGHT 1988). Column collapse can take place immediately after a single eruption, or during a series of closely timed explosions, such as those occur in many Vulcanian-style eruptions (NAIRN and SELF 1978). The resulting pyroclastic currents can be strik- ingly different. In a case of collapsing Plinian eruption plumes that may reach 30 km in height, large volumes of pyro- clasts can collapse into pyroclastic flows that radiate outward from the eruption centre. Since Plinian eruptions are gen- erally involve evolved and volatile-rich magma types, the resulting eruption clouds are charged with pumiceous pyro- clasts. Such eruption plumes can generate pumiceous pyroclast-dominated pyroclastic flows or ignimbrites. Small erup- tion plumes generated by Vulcanian-style eruptions (GOURGAUD et al. 2000) or vigorious lava fountain activity (BERTAGNINI et al. 1991; MASTROLORENZO et al. 1993; MARIANELLI et al. 1999; WOLFF and SUMNER 2000) intermittent with ongoing Strombolian-style eruptions commonly generate scoriaceous (more mafic) pyroclasts that can initiate high temperature pyroclastic flows comprising hot scoria and ash (GARDEWEG et al. 1998). Such currents termed as scoria- and-ash flows. Scoria-and-ash flows can also be generated by collapse of gradually accumulating scoriaceous deposits around an active vent because of over-steepening. Such an unstable pile of pyroclasts can perhaps collapse initiated a sub- sequent eruption. One major type of pyroclastic flow is associated with lava dome evolution on composite volcanoes. Lava domes are slow growing lava accumulations in the summit crater that slowly degas, and form a thickening crust. Such lava domes are hydrothermally altered and over time can become gravitationally unstable (Plate III, 3). When the lava dome reaches an unstable state, a relatively small volume, less energetic explosive eruption, commonly triggered by interaction between hot melt and hydrothermal systems, can lead to collapse of the lava dome (BOURDIER et al. 1997; SPARKS 1997; NAKADA et al. 1999; NA VARRO-OCHOA et al. 2002; BEHNCKE et al. 2003). Such dome-collapses then form a hot, avalanche-like current composed of large blocks derived from the lava dome and finer grained, generally ash-sized pyroclasts derived from the explosive eruption (ZOBIN et al. 2002; BEHNCKE et al. 2003). These currents are termed as block-and-ash flows. Dome collapse and edifice failure is known to reoccur and produce repeated episodes of block-and- ash flows and volcanic debris avalanches, as has been demonstrated on many volcanoes, e.g. Kamchatka, Russia (PONOMAREVA et al. 1998; BELOUSOV et al. 1999). Many block-and-ash flows are directly related to significant and ener- getic explosive eruptions and can be very hot. Such block-and-ash flows are also known as nuée ardentes or hot avalanch- 136es (WESTERCAMP 1987; BOURDIER et al. 1989; LAJOIE et al. 1989; BOUDON et al. 1993; TANGUY 1994; ABDURACHMAN et al. 2000; VOIGHT and DA VIS 2000; CARN et al. 2004; TANGUY 2004). Explosive eruption-triggered block-and-ash flows are commonly associated with laterally directed blasts that open up and destabilise the growing lava dome as occurred in the 1980 eruption of Mt St Helens (HOBLITT et al. 1981; VOIGHT 2000). However, edifice failure also can also provoke violent explosive activity, as was the case in the 1964 eruption of Shiveluch in Kamchatka, Russia (BELOUSOV 1995). A similar directed blast was formed dur- ing the 1956 eruption of Bezymianny volcano, Kamchatka, Russia (BELOUSOV 1996). This blast was generated by decompression of an intracrater dome and cryptodome that had formed during the preclimactic stage of the eruption (BELOUSOV 1996). The particle transportation and deposition- al processes are very similar for either type of pyroclastic flows. Many authors suggest that the resulting pyroclastic flow deposits do not always preserve key textural characteristics that allow us to interpret the transportation and depositional processes of pyroclastic flows (WILSON and HOUGHTON 2000). The trans- portation and depositional mechanism of pyro- clastic flows are therefore the subjects of ongo- ing research and debate. Two contrasting depo- sitional mechanisms are inferred (WILSON and HOUGHTON 2000) on the basis of the sedimentary textures of pyroclastic flow deposits;1) the progressive aggradation model (FISHER 1966, 1979, 1983; FISHER and SCHMINCKE 1984; BRANNEY and KOKELAAR 1992, 1994; FISHER and SCHMINCKE 1994; KOKELAAR and BRANNEY 1996) which postulates continuous deposition of an active flow base in the entire run out length of the flow. In this model the deposits therefore only give information about the base of the flow and the overall pyroclastic flow should be viewed as a high-density turbidity current from a physical point of view (KNELLER and BRANNEY 1995); 2) “en masse” freezing model which states that the entire flow freezes in a single moment (WRIGHT and WALKER 1981). The textural characteristics would therefore represent the textural characteristics of the flow itself. In this model, pyroclastic flows are viewed as cohesive debris flows (VALLANCE and SCOTT 1997; CAPRA and MACIAS 2000, 2002; CAPRA et al. 2002; LECOINTRE et al. 2002). Pyroclastic flow deposits Pyroclastic flow deposits result from all types of pyroclastic flows and predominantly consist of juvenile particles (FISHER and SCHMINCKE 1984; CAS and WRIGHT 1988; FISHER and SCHMINCKE 1994). Accidental lithics that have been picked up from the volcanic conduit and/or en route during deposition are relatively minor component in the resulting deposits. Juvenile particles are predominantly either pumice or scoria, depending on the source magma chemistry. Volcanic lithic fragments are especially common in block-and-as flows, where material is derived from the disrupted and collapsed lava dome (Plate III, 4). Pyroclastic flow deposits can also be associated with accretionary lapilli-bearing beds where the eruption was partially phreatomagmatic. Pyroclastic flow deposits are generally thick-bedded, matrix-rich and unsorted deposits that are commonly associat- ed with a thin basal ground surge and overlying ash cloud surge beds (Figure 6.4). The three major types of pyrcolastic flow deposits are common in textural point of view. Block-and-ash flow deposits are rich in non-to-moderately vesicu- lar lapilli and ash. Clasts in general are angular, and many of them show high temperature oxidation alteration. Larger lapilli and block size fragments are commonly jointed (Plate III, 5). The deposits are generally non-welded, however, 137 Figure 6.3. Types of pyroclastic flow generation according CAS and WRIGHT 1988: p. 106, fig. 5.11they commonly show evidence of high temperature depositional environment, such as large charcoalized logs, and as uniform magnetic fabric. Scoria-and-ash flow deposits are predominantly composed of scoria, many of them bed-flat- tened, especially in near vent locations (Plate IV, 1). Ignimbrite deposits (Plate IV, 2) are pumice rich and fine glassy shards comprise the matrix of the deposits. Gas escape pipes are common, especially in thick ignimbrite units (Plate IV, 3). Welding textures (Plate IV, 4) are common features in ignimbrite deposits and reflect of heat retantion of thick accu- mulations of pumiceous deposits. The base of a pyroclastic flow deposit is commonly marked by thin, cross bedded deposits from ground surges [(Plate IV, 5) immediately preceding the arrival of the main current body of the pyroclastic flow (SPARKS and WALKER 1973). The base of the pyroclastic flow deposit itself is richer in dense volcanic lithic fragments that may form lenses, or clast strings. The main flow deposit unsorted, and enrichment of pumiceous and/or scoriaceous fragments up-section is common (Plate V, 1). It is commonly covered by fine ash deposits that may be massive or slightly stratified. This unit is deposited from ash-clouds that emanate from passage of the main body of the pyroclastic flow current (FISHER et al. 1980a; VAZQUEZ and ORT 2006). In pumiceous pyroclastic flow deposits welding can occur (FREUNDT 1998; SPARKS et al. 1999; BEDDOE-STEPHENS and MILLWARD 2000; SPARKS et al. 2000). The high temperature of the individual pyroclasts can retain heat long enough that could lead to a partial melting of the fine-walled, low density clasts of evolved composition (WILSON and HILDRETH 2003). During the welding process, pumiceous and scoria- ceous pyroclasts become bed-flattened, and their vesicularity decreases due to the gradual collapse of vesicles. The more or less bed-oriented flattening of the pyroclasts and the associated plastic deformation of the entire unit lead to the formation of dense juvenile fragments of lenticu- lar shape, referred to as fiamme (Plate V, 2). The glassy shards of the matrix can sinter together to form a eutaxitic texture and produce the coherent, rock-like matrix of the fiamme-bearing succession. The weld- ing process can be very effective in thick pumiceous deposits, and could lead to the development of a single cooling unit over the entire succes- sion. The degree of welding is expressed in the grade of pyroclastic deposits (FREUNDT 1998; SUMNER and BRANNEY 2002). High grade pyroclastic flow deposits are rheomorphic, and can form lava-like units which are especially hard to distinguish from coherent lava bodies, particularly in ancient settings (BRANNEY et al. 1992; SMITH and COLE 1997; FREUNDT 1998; MUKHOPADHYAY et al. 2001; SUMNER and BRANNEY 2002; ALLEN 2004). Medium grade pyroclastic flow deposits are moderately welded (DUNCAN et al. 1999). Low grade pyroclastic flow deposits are non-welded (SZAKÁCS et al. 1998; YOKOYAMA 1999; EDGAR et al. 2002; THOURET et al. 2005). The distribution of pyroclastic flow deposits strongly depends on the morphology (Plate V, 3). Pyroclastic flows tend to follow topographic lows. Energetic pyroclastic flows, however, can flow uphill, flow over obstacles (Plate V, 4) (hundreds metres scale) (WOODS et al. 1998; LEGROS and KELFOUN 2000; GURIOLI et al. 2002) even up to hundreds of metres high (Taupo). Vegetation is commonly completely destroyed by pyroclastic flows (Plate V, 5). Valley pond deposits are generally thick units deposited from the axis of the pyroclastic flow currents (WILSON 1991; PITTARI et al. 2006) whereas veneer deposits are common along the valley margins and/or over obstacles. This lateral facies vari- ation is very common, and the resulting deposits are distinguishable (BROWN et al. 2003; PITTARI et al. 2006). The veneer deposits are commonly cross bedded, stratified, and finer grained than those deposited in the axis of the val- ley (FISHER et al. 1980b; WALKER et al. 1980, 1981; WILSON 2001; GIORDANO et al. 2002). From proximal to distal areas pyroclastic flow deposits also show some horizontal facies variations as a result of the loss of momentum by the cur- rent. Due to the physical properties of the pyroclastic currents, large but low density clasts can be transported far from their source. In contrast, dense volcanic lithic clasts decrease their grain size away from the vent. In near vent posi- tions, proximal, coarse lithic breccia units are common (DRUITT and SPARKS 1982; DRUITT 1985; DRUITT and BACON 1986; DRUITT 1995; WILSON 2001). Compositional zonation has been recognized in some large volume pumiceous pyroclastic flow deposits as a result of the ongoing and rapid emptying of the magma chamber (EDGAR et al. 2002; SUMNER and BRANNEY 2002). Magma mixing has also been recognized in large volume ignimbrites and may be a trig- gering mechanism for the eruption (BRIGGS et al. 1993; FREUNDT and SCHMINCKE 1995; HILDYARD et al. 2000; TROLL and SCHMINCKE 2002). Pyroclastic flow deposits because of the diverse mechanism of formation are very variable in volume. Scoria-and-ash flow deposits in general are less than 1 km 3 by volume. Large ignimbrites associated with caldera formation however, can form in excess of 1000 km 3 deposits (Plate VI, 1). The generally large volume of pyroclastic flow deposits means 138 Figure 6.4. Ideal, simplistic sections of major pyroclastic flow deposit types (after CAS and WRIGHT 1988). (a) block and ash flow deposit, (b) scoria-flow deposit, (c) pumice-flow deposit or ignimbritethat they are well preserved in the geological record. In long lived composite volcanic systems, pyroclastic flows can accumulate deposits to thicknesses of hundreds of metres, and therefore may play important role in the evolution of a sedimentary basin. Calderas and sedimentation associated with silicic volcanism Calderas are large depressions formed by subsidence driving voluminous volcanic eruptions (LIPMAN 2000). They are in general circular in map view, and can reach of depth of hundreds of metres (Plate VI, 2). They range up to several kilo- metres across and are commonly filled by caldera lakes (Plate VI, 3–4), and post-caldera eruptive deposits (Plate VI, 5) and associated reworked sediments. Caldera formation is generally associated with Plinian-style eruptions that form extensive Plinian pumiceous deposits and associated pyroclastic flows (ignimbrites) (e.g. Campanian Ignimbrite, Campi Flegrei, Italy) (Plate VII, 1). Pre-caldera successions are exposed in and the caldera walls and are overlain by syn-caldera volcaniclastic deposits and effusive products (LIPMAN et al. 1984). Very young calderas regardless to their well-preserved morphology do not give vital information about the subsidence processes and the root of the caldera itself. After forma- tion of a caldera, the central zone of the volcanic system becomes structurally unstable. Where the magmatic plumbing system remains active after caldera subsidence, the empty magma chamber being refilled by post-caldera magmas. The central zone of the volcanic structure may be raised (ACOCELLA et al. 2000; LINDSAY et al. 2001a; MASTURYONO et al. 2001). The uplift can be pushing significantly up the central part of the caldera (MASTURYONO et al. 2001) causing a sig- nificant geomorphic inversion over short period of time. Uplift can range in hundreds of metres scale and can cause a significant geomorphic inversion over relatively short period of time (thousands of years) (HULEN and NIELSON 1991; BATTAGLIA et al. 1999; NEWMAN et al. 2001). In the central part of the resurgent caldera system, caldera lake sediments can be raised to high stratigraphic positions, and post-caldera small-to-medium volume composite volcanoes can devel- op (KRUPP 1984; CHEN et al. 1995; TIBALDI and VEZZOLI 1998; MORAN-ZENTENO et al. 2004). The inversion results in the formation of resurgent dome emplacements. During the post-caldera formation, newly emplaced magma in the shallow feeding system can evolve and accumulate significant volume of volatiles, leading to further caldera forming eruptions. A resurgent caldera can experience repeated caldera subsidence, forming a complex architecture in the caldera volcano (ORSI et al. 1996; DI VITO et al. 1999). Such subsidence and resurgence can be cyclic, and lead to the development of nested, complex calderas (TIBALDI and VEZZOLI 2004; DE VITA et al. 2006). Although caldera forming eruptions are common in the geological record, few have been documented in historic times. These rare occurrences are generally small volumes and include the 1991 Pinatubo, Philippines (ROSI et al. 2001), 1968 Fernandia, Galapagos (ROWLAND 1996), 1912 Katmai, Alaska (HILDRETH and FIERSTEIN 2000), 1883 Krakatau, Indonesia (DEPLUS et al. 1995) or 1750–1790 Kilauea, Hawaii (SWANSON and CHRISTIANSEN 1973) eruptions. For example, the Pinatubo eruption only produced an approximately 2 km diameter caldera and formed up to 5 km 3 of dense rock equivalent (DRE) eruptive products. In the Cenozoic, however, there are many large volume caldera structures (WILSON et al. 1995) with deposits of hundreds of km 3 in volume. Such eruptions have also been consid- ered to have had climatic effects as well. Mega-calderas from the Cenozoic geologic record indicate eruption with potential devastating effects. Especially older, and resurgent, and therefore well-exposed volcanic successions of calderas indicate a typical caldera cycle (Figure 6.5) most of the calderas went through during their life (LIPMAN 1976; COLUCCI et al. 1991; LIPMAN et al. 1996; LIPMAN 1997). The sequence of events leading to caldera formation is marked by pre-caldera volcanism (COLUCCI et al. 1991). During this stage small magma batches erupt from initial accumulation of melts in shallow magma cham- bers. The gradual filling of the shallow magma chambers causes general uplift of the central region of the volcanic zone and formation of ring fractures and associated radial fractures, many of them filled with dykes. In this stage of the caldera development, magma may reach the surface and form a radial network of small-to-medium volume composite volcanoes, commonly with associated, long-lived effusive phases. Caldera subsidence takes place after favourable structural and magma chemical conditions in the central magma plumbing system trigger rapid eruption of large volumes of magmas. Although the resulting calderas are diverse in morphology, size, and the chemistry of the formed eruptive products, the generated calderas have a characteristic structure and morphology (Figure 6.6). The caldera rim is marked by a charac- teristic escarpment facing toward the central part of the caldera. This rim encircles both the structural boundary of the subsidence feature as well as the talus deposits formed by the collapse of the retreating topographic margin of the caldera. The inner topographic wall of a caldera forms a concave architecture and exposes pre-caldera successions, which may be covered by talus deposits formed by mass wasting in the rim. The collapse collar of the caldera is the region located between the structural caldera boundary and the inner topographic wall. In young calderas the collar slope angle (the line from the structural boundary to the topographic rim) is about 45 degrees but this can decrease to10-15 degrees in erod- ed structures. The structural caldera boundary is marked by bounding faults along which the caldera subsidence took place. These faults are ring-like and only exposed in older calderas where the caldera fill is removed and/or post-caldera 139resurgent uplift makes the deeper structure accessible. The calderas are generally filled with intra-caldera ash flow deposits and subsequent lacustrine successions. In the caldera topographic boundary, large blocks and caldera margin breccias may be associated with intra-caldera ignimbrites. The latter are commonly welded with multiple cooling units, and may form a strong, cap-like succession. The caldera floor is generally defined by the structural caldera floor, locat- ed well below the intra-caldera pyroclastic successions. The sub-caldera magma chamber (or remnant of it) can be found below the structural caldera floor in resurgent calderas, where the deep structures of the caldera can be in uplifted and exposed (LIPMAN 1984; JOHNSON et al. 2002). These magma chambers usually consist of plutonic rocks with strong hydrothermal alteration and common mineralization. As a final stage of the caldera cycle, small-to-medium volume vol- canic edifices may develop in the central part of the caldera. The most important evolutionary stage in the development of a caldera is its subsidence. Study the structural archi- tecture of the root zone of calderas, as well as detailed morphological analysis of young calderas aided by scaled model experiments, suggested few basic lines of subsidence styles (Figure 6.7) (ROCHE et al. 2000; WALTER and TROLL 2001; TROLL et al. 2002; HOLOHAN et al. 2005). The most common subsidence style is the piston or plate style of collapse of the caldera floor (ACOCELLA et al. 2000; ROCHE and DRUITT 2001; FOLCH and MARTI 2004). In this style the subsidence takes place along near vertical faults and the entire floor sub- sides as a plate. Trapdoor subsidence is an asymmetric type of subsidence, where an initial downsagging of the caldera floor takes place, followed by a plate-like col- 140 Figure 6.5. Caldera cycle after LIPMAN 2000: p. 648, fig. 2] Figure 6.6. Morphological and structural elements of a caldera after LIPMAN 2000: p. 649, fig. 3lapse (MAHOOD and HILDRETH 1983; BERESFORD and COLE 2000; MILNER et al. 2002; KENNEDY et al. 2004; RAMELOW et al. 2006). Pure downsagging is apparently rare (ORT 1993). Piecemeal fault controlled subsidence is very common in many calderas, as a result of differential subsidence along many well defined structural weakness zones (ROSI et al. 1996; MOORE and KOKELAAR 1997, 1998; MILNER et al. 2002). Such subsidence can result differently subsided blocks in the central part of the caldera floor. Any of the above types of caldera subsidence can lead to chaotic subsidence which is also probably the most com- mon styles in complex and large mega-calderas. Funnel-shape subsidence is an end mem- ber of subsidence types, and probably only associated with small calderas. The internal structure of calderas can be studied when the resurgence of the caldera floor uplift the cen- tral part of the caldera as it is the case in the Valles Caldera in Texas, one of the most well- studied caldera in Earth (Plate VII, 2). Well-defined calderas and associated ignimbrite sheets are well known and well described in many Cenozoic calderas. In general the identification of large volume and extensive ignimbrite sheets are therefore often used to infer the location of associated calderas. However, in many regions very extensive ignimbrite sheets (Plate VII, 3) had no apparent association with caldera structures, such as the widespread and large vol- ume Cenozoic ignimbrites in the Andes (LEBTI et al. 2006) or in Armenia. Such silicic ignimbrites are commonly referred as ignimbrite shields (DE SILVA 1989). Only in recent times have researchers recognised that these ignimbrite shields are also associated with caldera of structural elements (ORT 1993; LINDSAY et al. 2001b; RICHARDS and VILLENEUVE 2002). Identification of caldera structures, and separating their structural elements from regional tectonic structures is difficult in areas where ignimbrite-forming eruptions are associated with regional extension and volcanic products accumulate in a volcano-tectonic structure, such as the Taupo Volcanic Zone of New Zealand (WILSON et al. 1995). Similar problems hinder the identification of the source of extensive Neogene basin-wide ignimbrite sheets in the Miocene Pannonian Basin (PANTÓ 1963; SZABÓ et al. 1992; CAPACCIONI et al. 1995; PÓKA et al. 1998; SZAKÁCS et al. 1998). Similar volcano-tectonic graben structures are inferred to contour a chain of large calderas in the Tokaj Mountains in Hungary. However, more detailed studies still need to be done confirm this interpretation of the preserved volcanic units (MOLNÁR and ZELENKA 1995; SEGHEDI et al. 1998; BAJNÓCZI et al. 2000; PÉCSKAY and MOLNÁR 2002; MARTIN et al. 2003). In the Taupo Volcanic Zone, it also has been recently recognized, that certain ignimbrite sheets can be associated with previously not recognized eruptive centres, and growing number of hard to recognize calderas have been identified (NAIRN et al. 1994; WILSON et al. 1995; BERESFORD and COLE 2000; WILSON 2001; MILNER et al. 2002, 2003). The size of calderas is very diverse, but many evidences in the geological record suggest that even the larger calderas can form in a relatively short period of time. Since we don’t have experience of such eruptions in our history, the vol- canic hazard implication of caldera eruptions is poorly understood. They are however, potentially the largest and most hazardous eruptions on Earth. Such eruptions commonly are considered to also be climate-forcing and even effect evo- lution (FEDELE et al. 2002; OPPENHEIMER 2003a, b; MANVILLE and WILSON 2004; SELF 2006). However, because of the more evolved composition of magmas known to be involved in caldera formation, the climatic effect of such eruptions could be smaller than those medium-volume, but more mafic composition volcanoes that produce larger volumes of sul- phur aerosols, known to have a large impact on the Earth’s thermal balance. Post volcanic hydrothermal activity Post-volcanic hydrothermal activity of polygenetic volcanoes, especially silicic ones are is very diverse and play major role in mineralization. In caldera systems, the shallow magma chamber still can provide heat source effective enough, to heat ground water, and produce solfatara, mofetta or geyser activity long time over eruption, and/or inter-erup- tion periods. Solfatara (Plate VII, 4) is a dominantly sulphur producing vent named after Solfatara, a small post-caldera tuff ring in the Campanian Field in Italy. Mofetta is a CO 2 producing vent, commonly associated with solfatara fields. Major solfatara fields in silicic volcanic systems are commonly accompanied with small hydroexplosion sites (Plate VII, 5). Due to the elevated heat of the shallow magma chamber provide to the ground water, overpressurised hydrothermal systems, time to time can outburst into hydroexplosions, and may form few tens of metres wide craters surrounded by dm-to-m thick chaotic breccias (Plate VIII, 1) dominated by hydrothermally altered volcanic lithic fragments. Such 141 Figure 6.7. Styles of caldera subsidence after LIPMAN 2000: p. 654, fig. 6hydroexplosion (phreatic explosions) sites are commonly forming extensive (few km 2 area) fields (Plate VIII, 2), a per- fect sites for mineralization. Explosion craters are commonly filled with thermal water and can form reasonable sized volcanic craters similar in architecture to maars (Plate VIII, 3). Water level of such craters can change dramatically (Plate VIII, 4). Hydrothermal fields commonly have geysers, and associated outflow springs rich in minerals. Such mineral- rich waters may accumulate laminated lacustrine beds in shallow hot water pools in the thermal areas (Plate VIII, 5). In ancient settings similar shallow lacustrine systems maybe preserved in a form of laminated fine grained lacustrine sedi- ments composed of angular volcanic lithics as well asclay minerals such as deposits in the Miocene Tokaj Mts in NE- Hungary (Plate VIII, 6). Hydroexplosions can form over ignimbrite sheets. Ignimbrite can retain heat long enough after deposition and generate overheated steam that may disrupt in an explosion the overlying ignimbrite sheets. Such erup- tions are documented from the Mount St Helens. The resulting deposits and crater morphology is similar to those formed by phreatomagmatic eruptions, however, there is no juvenile magma involvement in such hydroexplosions, and therefore the deposits will be only consist of accidental lithics derived from the ignimbrite units. References ABDURACHMAN, E. K., BOURDIER, J. 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A., URSUA, M., VELASCO, J., ALATORRE, E. and SANTIAGO, H. 2002: Overview of the 1997–2000 activity of Volcan de Colima, Mexico. — Journal of Volcanology and Geothermal Research 117 (1–2), pp. 1–19. 149150 Chapter 6 Location map 1 — Hawaii, Kilauea, USA 2 — Tenerife, Spain 3 — Al Haruj, Libya 4 — Eastern Australia 5 — Karoo, South Africa 6 — Ceboruco, Mexico 7 — Hekla, Iceland 8 — Askja, Iceland 9 — East African, Rift 10 — Mayon Phillippines 11 — Ararat, Turkey 12 — Unzen, Japan 13 — Campi Flegrei, Italy 14 — Pinatubo, Phillippines 15 — Fernandia, Galapagos 16 — Katmai, Alaska, USA 17 — Krakatau, Indonesia 18 — Valles Caldera, Texas, USA 19 — Armenia 20 — Pannonian Basin 21 — Tokaj Mts, Hungary 22 — Vulcano, Italy 23 — Santa Catarina Mts, Sonora, Mexico 24 — Cappadocia, Turkey 25 — Mt. Peleé, Martinique 26 — Gran Canaria, Spain 27 — Mendoza, Argentina 28 — Ischia, Italy 29 — Usu, Japan 30 — Rininahue, Chile 31 — Calbuco, Chile 32 — Osorno, Chile 33 — Shiveluch, Russia 34 — Bezymianny, Russia 35 — Karymsky, Russia 36 — Toyo Caldera, Japan 37 — Hokkaido-Komagatake, Japan 38 — Taranaki Volcano, New Zealand 39 — Lake Taupo, New Zealand 40 — Ngauruhoe Volcano, New Zealand 41 — White Island, New Zealand 42 — Tarawera, New Zealand 43 — Crater Lake, Oregon, USA 44 — Mt St Helens, Washington, USA151 Chapter 6 Plate I 1. Shield volcano of the Tenerife Island (in horizon) having broad base composed of thick succession of lava flows and subordinate pyro- clastic rock units similar to the succession in the fore- ground of Gran Canaria, Spain (photo by S. J. Cronin). 2. Hyaloclastite successions of the old shield edifice of Tenerife Island, Canary Islands, Spain 3. Medium size Pliocene intracontinental shield vol- cano from the Al Haruj al Abyad volcanic field, Central Libya. 4. Low-relief, Pliocene medium size lava shields forming an amalgamated network of volcanic land- forms along the Snake River Plain in Idaho. 5. Lava delta in cross sec- tional view along the Snake River valley, Idaho.152 Chapter 6 Plate II 1. Crater of Volcan Ceboruco in the western Transmexican Volcanic Belt. The crater of Ceboruco is occupied by small volume lava coulees and domes (centre of view). 2. Satellite vents on the flank of the Santa Catarina shield volcano, in the Pinacate Volcanic Field in Sonora, Mexico. 3. Lava dome complex on top of an arc stratovolcano in Northern Chile forming a complex summit morphology of the stratocone. 4. Truncated summit region and morphology of Calbuco Volcano in Southern Chile a result of multiple collapse events. 5. Strongly modified and glaciated summit of a southern Chilean stratovolcano as a result of strong external erosional forces and mul- tiple collapse events.153 Chapter 6 Plate III 1. Conical shape of Taranaki stratovolcano, New Zealand. 2. Summit region of Osorno in Southern Chile with a young and fresh scoria domi- nated cone evolving over a truncated possi- ble collapsed summit region of the vol- cano. 3. Hydrothermally altered “ready-to-col- lapse” architecture of the Hokkaido- Komagatake lava dome in Hokkaido, Japan 4. Cross section of block-and-ash flow deposits from the 1973–75 eruption of the Ngauruhoe volcano, New Zealand. 5. Prismatically jointed lava block in block and ash flow deposit from the 1980 eruption of Mt St Helens (photo by S. J. Cronin).154 Chapter 6 Plate IV 1. Scoria-and-ash flow deposit from the 1979 eruption of Ngarahue volcano, New Zealand. Note the levees formed by the channelised scoria-and-ash flow deposit. 2. Ignimbrite deposits from the Central Anatolian Ignimbrite Province, Zelve Valley, Cappadocia, Turkey. 3. Gas escape pipes preserved as pillars in ignimbrite deposits of the Mazama eruption, Crater Lake (photo by S. J. Cronin). 4. Welded ignimbrite unit of the Crater Lake, Oregon (photo by S. J. Cronin). 5. Ground surge deposit of Mt Pelee 1904 eruption exhibit dune bedded structure (photo S. J. Cronin). 155 Chapter 6 Plate V 1. Main pyroclastic flow unit in Miocene non-welded ign- imbrite with elutriation pipes from the Tokaj Mtns, NE Hungary. 2. Fiamme in welded ign- imbrite from Gran Canaria (photo by S. J. Cronin). 3. Valley filling ignimbrite pinching out in the palaeo- valley margin (Mendoza, Argentina). 4. Ignimbrite deposits (white blanket on the flank of the cone) over scoria cones in southern Tenerife (Montana Pelada tuff ring) indicating high inertia of the pyroclastic flow running over tens of metres high obstacles. 5. Fallen trees on small hill side the 1980 pyroclastic surge and flow run over at Mt St. Helens (photo by S. J. Cronin).156 Chapter 6 Plate VI 1. Extensive sheet-like ign- imbrite plateau in the Altiplano, Northern Chile. 2. Taupo caldera with Lake Taupo, New Zealand. Note the steep escarpment along the caldera lake shore formed by ignimbrite successions. 3. Toyo caldera lake in MrSID satellite immage, Hokkaido, Japan. The islands in the lake (black) are resurgent domes. 4. Karymsky caldera in Kamchatka on oblique satel- lite image (Google Earth). Note the intra-caldera Surt- seyan-style tuff cone in the shoreline of the caldera lake formed in 1996. The tuff cone is eroded by now, and only a pyroclastic mound left on the lake floor. 5. Crater Lake caldera with Wizard Island in its centre, as a post caldera scoria cone (photo by S. J. Cronin)157 Chapter 6 Plate VII 1. Oblique view of a satellite image of the Campi Flegrei, a large caldera system produced the Campanian Ignimbrite. The caldera structure is occupied by small post caldera tuff rings, and scoria cones. 2. Resurgent caldera structure of Ischia in Italy on oblique satellite image (GoogleEarth). 3. Extensive ignimbrite shield in Northern Chile. The yellowish landscape on the entire view is an extensive ignimbrite shield formed in the Quaternary. 4. Solfatara field in the inner crater wall of Vulcano, Lipari Island, Italy. 5. Deep hydrothermal explosion crater in New Zealand.158 Chapter 6 Plate VIII 1. Hydrothermal explosion breccia from New Zealand. 2. Hydrothermal explosion crater field in New Zealand. 3. Maar-like explosion craters around Tarawera, New Zealand. 4. Water level changes represented in narrow benches in the shoreline of an explosion crater near Tarawera, New Zealand. 5. Thermal water pools in thermal fields in the New Zealand. 6. Miocene hydrothermal deposits associated with a thermal area developed over Miocene silicic pyroclastic successions of the Tokaj Mtns, NE-Hungary.Volcanoes produce large amounts of brecciated materials that contribute to the construction of a stratocone, and accu- mulate around the volcanic edifice to form a ring plain or volcaniclastic apron. Large volume volcanic edifices (i.e. >100 km 3 ) form over long periods of time (e.g. hundreds of thousands of years). During periods of quiescence, volcanoes can be effected by external forces (e.g. climatic agents) that continuously remove parts of the volcanic edifice and redistrib- ute volcanic debris to the surrounding terrains (MATHISEN and MCPHERSON 1991, RAMPINO 1991, SCHMINCKE 2004). Redistribution of primary volcanic deposits can take place immediately after, or during the course of an eruption, when volcanic processes rapidly initiate secondary depositional mechanisms. Volcano-sedimentary facies evolve almost con- tinuously between primary and secondary volcanic deposits, in almost any type of volcanic system. Secondary volcanic processes that generate large volumes of sediments, such as lahars, may also be initiated by external forces such as heavy rainfall and storms, a common occurrence in tropical environments (SMITH and LOWE 1991, FISHER and SCHMINCKE 1994). Flank instabilities and sector-collapses have also been recognized (VOIGHT and ELSWORTH 1997, VOIGHT 2000) as key mechanisms leading to the catastrophic emplacement of volcano-sedimentary successions, especially around polygenetic, long-lived strato-volcanoes. Cone collapses generate volcanic debris avalanches that often transform into debris flows in distal reaches of the ring plain (SMITH and LOWE 1991, UI et al. 2000). Long-lived composite volcanoes usually reach an advanced stage of growth before flank collapse occur, resulting in exceptionally large volume deposits (tens of cubic kilo- metres) made of chaotic volcaniclastic successions accumulated around the central volcanic edifice (SIEBERT 1984, BEGET and KIENLE 1992, NEHLIG et al. 2001). Such catastrophic volcanic collapses are commonly triggered by moderately explosive eruptions (e.g. Shiveluch Volcano, Kamchatka; PONOMAREV A et al. 1998). However, such collapses could also take place dur- ing periods of quiescence, when mechanical failure of unstable portions of the edifice are suddenly triggered by hydrother- mal circulations, seismo-tectonic events or dyke emplacement (VA N WYK DE VRIES and BORGIA 1996, VA N WYK DE VRIES 1998, VA N WYK DE VRIES and MATELA 1998, VOIGHT 2000, TIBALDI 2001, TIBALDI et al. 2003). In fact, gradual spreading of an unstable volcanic edifice over long periods of time (thousands of years) may also truncate, redistribute and modify origi- nal primary volcanic successions with no single catastrophic event involved (VA N WYK DE VRIES 1998, LAGMAY et al. 2000, CECCHI et al. 2004). Studying volcanic successions accumulated around stratovolcanoes may help to identify generating processes of lahars and volcanic debris avalanches that contributed to the construction of the ring plain. In both cases, the key question is, what the triggering mechanisms are, and how closely that triggering mechanisms are linked to eruptive activity (e.g.: was a volcanic eruption involved in the formation of lahars and volcanic debris avalanches or not?). Volcanic sediment redistribution can also take place without any sudden, catastrophic mass-wasting. Instead, sus- tained erosion of the original volcanic edifice may lead to the development of specific erosional landforms, and the accu- mulation of volcaniclastic sediments in the surrounding sedimentary basins. Effects of (or impact of) long term erosion on volcanic landforms are poorly understood, and the link between the original edifice and the resulting accumulated vol- caniclastic sediments may be difficult to establish with certainty in older volcanic terrains. Volcanic ring plain Volcanic ring plain was introduced in the literature from studies of central New Zealand volcanoes (Table I, 1) where volcaniclastic sequences have been mapped in detail, and key facies changes observed between the central volcanic edi- fice and the surrounding low lands have been properly recorded (NEALL 1975, PALMER and NEALL 1991, PALMER et al. 159 Chapter 71993, CRONIN and NEALL 1997, LECOINTRE et al. 1998, DONOGHUE and NEALL 2001). A volcanic ring plain is defined as the circular area that surrounds a centrally constructed volcanic edifice (usual- ly a stratovolcano) (Figure 7.1). Ring plains are therefore sedimen- tary basins where both primary and secondary volcaniclastic prod- ucts accumulate (PALMER et al. 1993). During periods of volcanic activity, the ring plain is progressively built with the deposition of primary volcanic products (CRONIN et al. 1996b, 1996c). Layers of pyroclastic debris (ash, lapilli) mantle the volcaniclastic apron (Table I, 2), preferentially accumulating in sectors of the ring plain where dominant winds control the direction of the eruption cloud (NEALL 1975, PALMER and NEALL 1991, ALLOWAY et al. 1995). Such fall deposits can be used successfully as marker beds for cal- ibrating key lithostratigraphic sections, with an aim to establish the detailed record of volcanic and mass wasting events that affected a centrally constructed volcanic edifice and its surrounding ring plain. Because the wind direction change over time, each individ- ual fall deposit may accumulate in a slightly different area of the volcanic ring plain, as demonstrated around the main central North Island volcanoes of New Zealand (FROGGATT and LOWE 1990, DONOGHUE and NEALL 1996, LOWE et al. 1998, DONOGHUE et al. 1999, SHANE 2000) and elsewhere (SCHMINCKE and VAN DEN BOGAARD 1991, ANDREASTUTI et al. 2000, LEGROS 2001). In extreme cases, fluctuating wind directions during successive erup- tion phases can lead to the accumulation of tephra beds that are not overlying each other (Figure 7.2). If magma sources are geochem- 160 Figure 7.1. Diagram shows major facies associations around an active composite volcano (after MATHISEN and MCPHERSON 1991: pp. 29, fig. 3) Figure 7.2. Tephra successions of pyroclastic fall unit form a complex stratigraphy in the ring plain around the Tonagariro volcano in New Zealand (after DONOUGH and NEILL 1996). A, B and C shows different fall unit thicknesses as a result of different wind direction during their formationically very similar, the correct identification of specific fall tephra units could be challenging, and thus, leading to incor- rect tephrostratigraphic interpretations. Tephra accumulating on the ring plain are commonly intercalated with other vol- caniclastic deposits resulting from erosional and/or major destructional processes that affected the volcanic edifice (PALMER et al. 1993, LECOINTRE et al. 1998, LECOINTRE et al. 2004). The ring plain is also a place where distal vol- canogenic deposits sourced from pyroclastic flows and primary lahars may accumulate and mix (PALMER 1991, CRONIN et al. 1997b, LECOINTRE et al. 1998). In ancient volcanic terrains, volcanic deposits commonly represent lithostratigraph- ic successions formed in a sedimentary basin around volcanoes, such as ring plain associations (BREITKREUZ 1991, SPALLETTI and DALLASALDA 1996, AYDAR 1998, CORCORAN et al. 1999, CUMMINGS et al. 2000). Because of the incom- plete preservation of primary deposits on the slopes of the cone, ring plain volcaniclastic sequences could reveal a more detailed chronostratigraphic record that include discrete volcano-sedimentary events of the past. Therefore, the study of these distal volcanogenic accumulations is an essential step to properly reconstruct the eruptive history of a stratovol- cano. To understand the volcanic architecture of a complex, ancient volcanic system, it is thus very important to inter- pret volcanic terrains in the light of understanding the depositional processes that contributed to the construction of the volcanic ring plain. Lahars Lahar is an Indonesian term defining a whole spectrum of laterally spreading volcanic mass flows such as debris flows, transitional flows or hyperconcentrated streamflows that originated from a volcano by any genetic processes (SMITH and LOWE 1991, FISHER and SCHMINCKE 1994, LA VIGNE and THOURET 2000, VALLANCE 2000). Lahar therefore is a very broad type of gravity-driven volcaniclastic current that can produce many different types of deposits (Figure 7.3). Lahar as a term should be used to describe the process, not the deposit (SMITH and LOWE 1991). Lahars deposits there- fore have to be carefully examined in order to determine the physical process(es) responsible for their emplacement. Lahars can be generated directly by a volcanic erup- tion (WAITT et al. 1983, MOTHES et al. 1998) or by sudden rainfall that quickly remobilise volcanic debris on the flanks of an active volcano (HODGSON and MANVILLE 1999, VA N WESTEN and DAAG 2005). Lahars can also be trig- gered by catastrophic pre- cipitations induced by hur- ricanes over a dormant vol- cano such as the Casita in 1998 in Nicaragua (SCOTT et al. 2005). Deposits resulting from these two distinct processes could look very similar to each other by their texture, bed- ding characteristics or composition. Lahars start often as minor debris flows that may transform into debris flows incorporating large volume of volcanic and non-vol- canic debris en-route (SMITH and LOWE 1991, VALLANCE 2000). Lahars commonly reach the ring plain (Table I, 3) and form extensive sheet-like deposits intercalated by other primary deposits such as ash layers (PALMER et al. 1993). Lahars represent a very significant hazard in volcanoes located in tropical countries such as Mt Merapi in Indonesia (LA VIGNE et al. 2000, LA VIGNE and THOURET 2003), or Mt Pinatubo in the Philippines (CHOROWICZ et al. 1997, VA N WESTEN and DAAG 2005, CARRANZA and CASTRO 2006). Ancient lahar deposits interbedded with palaeosols and loesses may also provide clues on palaeoclimatic conditions of the region where the volcano erupted. Lahars are also commonly pro- duced by active volcanoes that have craters occupied by a lake, such as Mt Ruapehu in New Zealand (CRONIN et al. 1996a, LECOINTRE et al. 1998, MANVILLE et al. 1998, CRONIN et al. 1999). Eruption through a crater lake may suddenly expel large amounts of water on the steep, tephra-covered, upper slopes of the cone. This violent outpouring of hot and 161 Figure 7.3. Relationship between lahar types, their deposits, and resulting theoretical sedimentary structures after SMITH and LOWE 1991: p. 60, fig. 1) often acidic waters initiate debris flows that by bulking process, will incorporate large volume of extra sediment (sand, boulders) on the way down to the ring plain (MASTIN and WITTER 2000). Eruptions through crater lakes are considered to be dangerous, a common hazardous situation that affect exposed populations on volcanic islands such as Ambae in Vanuatu (NÉMETH et al. 2006a). Lahars can also be initiated by sudden melting of a glacier, or snow- and ice-cap (Table I, 4 and 5), that often cover the summit region of a high composite volcano (MANVILLE et al. 2000) and may lead to the rapid formation of “snow slurry lahars” (CRONIN et al. 1996a). Such processes can be triggered by a growing lava dome beneath the summit of the volcano. An elevated heat gradient, due to hydrothermal and magmatic fluids in the upper flank of the edifice, may lead to the sudden melting of the snow-and-ice cap. Such catastrophic process is unfortunately very common on the highly elevated volcanoes of the Andean volcanic arc, and is responsible for devastating mudflow inundations in surrounding valleys (BRANNEY and GILBERT 1995, STERN 2004). Lahars initiated from the sudden melting of an ice cap are able to carry large ice blocks (tens of m 3 ) that can be deposited at great distances from source, downstream the river catchment. After melting of the ice, voids may form in the sandy deposit that once collapsed, will be filled with chaotic debris (CRONIN et al. 1996a). Lahars also carry large volumes of eroded soil, vegetation, and truck- or house-sized debris from infrastructure destroyed on their path. Lahars are also commonly able to run over small obstacles (tens of meters high), flow uphill over small morpho- logical barriers, and can reach long distances (tens of kilometres) from their source (CARRASCONUNEZ et al. 1993, KOHLBECK et al. 1994, MOTHES et al. 1998, HALL et al. 1999). Such lahars are therefore a real threat to unaware pop- ulations that live or travel in areas generally considered to be safe from a distant — and often dormant — volcano. Large volume deposits resulting from these catastrophic lahars may represent a significant proportion of the sedi- mentary basin fill, even at far distances from the source volcano. This observation implies that identification of laharic deposits in ancient settings should be conducted very carefully, in order to determine the exact source loca- tion of the studied diamictons (WALTON and PALMER 1988, SPALLETTI and DALLASALDA 1996, KARÁTSON and NÉMETH 2001, NEHLIG et al. 2001, ENCINAS et al. 2006). Lahars generally flow down the valley network radiating around the volcano, and accumulate deposits in major basinal areas of the ring plain. The sudden input of large volume of vol- canic debris in existing fluvial networks may significantly alter the course of the main river and its secondary trib- utaries (Table II, 1). Therefore, modelling and prediction of lahar inundation areas is an essential but complex task, especially in volcanic fields where lahars can occur frequently without too much warning (e.g. in tropical volcanic islands). The generation of lahars and the volume of resulting deposits could vary hugely. After the Pinatubo 1991 erup- tion, tephra covering vast expenses of bare surfaces was quickly remobilized into lahars after heavy rain falls, with no volcanic eruption involved (GRAN and MONTGOMERY 2005, VA N WESTEN and DAAG 2005, CARRANZA and CASTRO 2006). Over a 3 years period, 30% of the total tephra volume produced by the 1991 eruption have been remobilized and accumulated in stream valleys and onto the Pinatubo ring plain (TORRES et al. 2004). Lahar deposits can be sub- sequently recolonized by vegetation once soil development occur and form a stable landscape (OBA et al. 2004). Another famous lahar, which occurred on the Nevado de Ruiz in Columbia in 1985, was initiated by medium vol- ume hot pyroclastic flows that interacted with the ice- and snow-cap (KOHLBECK et al. 1994, PIERSON and JANDA 1994). Initial melt water mixed with the ejected volcanic debris, and produced a small volume lahar that evolved into a series of devastating, widespread lahar waves (0.1 km 3 ), killing c. 23,000 people. The lahars reached zones located 100 km from the source, and modified significantly the drainage network of the area. Nevado del Ruiz-type lahars are commonly triggered by pyroclastic flow-generating eruptions. The resulting mass flows cover the lower flanks of the volcano and the surrounding ring plain, where a complex network of intercalated and inter-fingering primary and secondary volcaniclastic sequences form a fan-like sedimentary cover (Table II, 2). This succession commonly referred as block-and-ash fan, has been described initially at the Pico de Orizaba, Mexico (SIEBE et al. 1993). Laharic diamictons in ancient volcanic settings are common. Many of them correspond to very large volume of deposits produced by a single mass wasting event such as the Osceola Mudflow from Mount Rainer, in the Cascades Range, Washington State, USA (VALLANCE and SCOTT 1997). The giant, clay-rich lahar initiated catastrophically from a fluid-saturated debris avalanche, deposited 3.8 km 3 of volcaniclastic sediment over a distance >120 kilometres (VALLANCE and SCOTT 1997). Similar volume of lahar deposits are also known in South America, for instance around Mt Cotopaxi in Equador (AGUILERA et al. 2004). The Chillos Valley lahar has also been triggered by pyroclastic flow-gen- erating eruptions, and reached a distance of 300 km from their source. It filled radiating valleys with sediment accumu- lations up to 200 meter deep (MOTHES et al. 1998). The preservation potential of lahar deposits –especially for debris flows- is in general good. Lahar are quickly chan- nelled in valleys before reaching the ring plain, and their resulting deposits may contain interbedded layers hardened by dewatering processes, and also, many large lithic blocks that reinforce their broad compaction. Therefore, confined lahars tend to “fossilize” palaeo-valley networks around a volcano. 162From volcanic debris flows to hyperconcentrated (flood) streamflows Lahar, as a volcaniclastic mass flow initiated on the flank a volcano, can evolve quickly during its course downstream a river channel. As a result, very different types of deposits can form, showing a broad horizontal facies variations along the stream valleys (LA VIGNE and SUWA 2004). If lahars are initiated on a volcanic terrain covered by abundant, loose, gen- erally fine-grained sediment (e.g. ash), they can quickly transform from hyperconcentrated streamflows (HcFs) to debris flows (DFs), depending on the amount and size of particles the lahar current may pick up en route (PIERSON and SCOTT 1985, SMITH and LOWE 1991, CRONIN et al. 1997a, VALLANCE 2000). Further downstream, the lahar may become more diluted, as a water-rich hyperconcentrated flow that could quickly change into normal flood, generating sediment-laden streamflow deposits (SMITH and LOWE 1991, VALLANCE 2000). Volcanic debris avalanches initiated by sudden collapse of a volcanic edifice tend also to transform rapidly into debris flows, then into hyperconcentrated flows and normal floods downstream, as the mass flow hits water stored in lower inundation zones (PIERSON and SCOTT 1985, SMITH and LOWE 1991, CRONIN et al. 2000). During lahar emplacement, significant erosion can affect the bed rock units, especially when the lahar wave moves over erodable terrains (e.g. loose tephra cover) (SMITH and LOWE 1991). Undercutting of channel walls may also cause significant erosion and incorporation of sediment into the active lahar current (SMITH and LOWE 1991). This erosional effect can be significant during both debris flow and hyperconcentrated flow phases of the lahar current. As previously mentioned, the lahar current may incorporate downstream large volumes of sediments (bulk- ing), a process that is responsible for the transition from normal flood flow events to hyperconcentrated flows and debris flows (SMITH and LOWE 1991). Such mechanism is observed specifically during the waxing and waning phases of the flow. Sediments picked up during the passage of the lahar waves tend to be segregated into distinct beds rich in exotic lithologies. Large volume debris flows can be loaded with volcanic and non-volcanic lithic fragments, colluvium, tree debris as well as glacial drifts. The waxing phase of such lahar currents can be very erosive (SMITH and LOWE 1991). The fine sediment-rich tail of the lahar is generally less erosive that the front wave, facilitating the deposition of sediments (SMITH and LOWE 1991). The waning stage of the lahar current is in general water-rich, of smaller amplitude than the frontal head, but still rather erosive. At this late stage of evolution, it incises into the previously deposited deposits and can create complex sedimentary features in cross sectional view through the stream channels (SMITH and LOWE 1991). Lahars go through significant particle segregation processes (Figure 7.4) due to density and particle concentration variations throughout the entire current (PIERSON and SCOTT 1985, SMITH and LOWE 1991). Light, pumiceous clasts quick- ly reach the top of the lahar current. In a similar way, low density solids also tend to pop up on the surface of the current due to pore water escape in the body of the flow. Such low density particles commonly form rafts on the sur- face of the current, and can travel together over long distances (SMITH and LOWE 1991). Due to the viscous nature of the lahar current, its bottom part moves more slowly then its top part. This differentiation leads to the development of typical vertical flow profiles (for veloc- ity and particle concentration) (SMITH and LOWE 1991). The low density particles therefore migrate upward, then to the front of the lahar wave. Conversely, high density particles tend to settle quicker. This particle segregation process results in a typical normal grading texture for the lahar deposits. However, kinetic sieving also occurs in a vibrating or sheared grain flow, as a result of both percolation and fluid expulsion. This process, induces downward movements of small parti- cles, progressively replacing larger clasts (SMITH and LOWE 1991). As a consequence, large particles tend to reach the surface of the flow, and drift progressively towards its lateral margins, depositing lenses of well sorted sediments. During its migration downstream a river catchment, a lahar will gradually incorporate more water from the stream channel in amounts large enough to change the physical properties of the flow and drop its capacity to carry sediments (PIERSON and SCOTT 1985, THOURET et al. 1998, CRONIN et al. 1999, LA VIGNE et al. 2000, LA VIGNE and SUWA 2004). This process is important 163 Figure 7.4. Particle-size segregation through a laharic current after VALLACE 2000: p. 607, fig. 3only for those lahars that are small enough in comparison to the water course they enter. The rheology and physical behaviour of large volume lahars are generally not affected significantly by this mechanism (LA VIGNE and THOURET 2000). Small volume clay-poor lahars (i.e. “non-cohesive” lahars; SCOTT 1988, SCOTT et al. 1995) tend to have an open framework, sandy matrix, and can be effected by dilution from water (MOTHES et al. 1998). By contrast, clay-rich lahars (or “cohesive” lahars) are harder to dilute due to the buffering effect of clay particles in the matrix. Lahars that reach the ring plain and follow major tributaries push the stream water ahead of the front wave, and then gradually mix. Water ingestion into the lahar body leads to a drop in its capacity to carry large clasts, a process that leads to the formation of lag-like lenses of cobbles and boulders. The textural characteristics of the resulting deposits depend on the type of lahar flow (Figure 7.5). Debris flows are generally poorly sorted, massive, and non-stratified (Table II, 3). By contrast, hyperconcentrated stream flow deposits are better sorted, and show a faint internal stratification (Table II, 4). Textures of hyperconcentrated streamflow deposits may also reveal clast alignments with compositional variations among clast groups, leading to a graded vertical distribu- tion of coarse particles. Some of these clast groups show evidences of internal abrasion (cataclasis), and tend to be asso- ciated with more stratified units. Distinguishing lahar-originated deposits (Table II, 5 and 6) from those non-welded fine grained and matrix-rich pyro- clastic flow deposits is sometimes difficult (BRANTLEY and WAITT 1988, WAYTHOMAS 1999). However, detailed studies of lateral facies variations can be helpful in order to distinguish the textural characteristics of non-welded pyroclastic flow deposits (juvenile particle content, vesicle-free, and non-indurated matrix) from those of lahar deposits. Since pyroclas- tic flows are hot during emplacement, their deposits commonly contain charcoals as a result of burnt vegetation (SCOTT and GLASSPOOL 2005), but a distinctively rare occurrence in lahar deposits (Table III, 1). However, pyroclastic flows can also hit major water courses, snow-and-ice caps, and be partially diluted by heavy rain, thus initiating hot lahars (Table III, 2 and 3). When lahars are generated in a valley previously filled by a pyroclastic flow, the mass flow will remobilise the pumiceous material that will be re-deposited in more distal areas of the ring plain (Table III, 4). In this case, lahar deposits can incorporate charcoals that have been originally contained in a pyroclastic flow deposit channelled further upstream in the valley (Table III, 5). Pyroclastic flow deposits are usually hot enough to develop magnetically oriented clasts, and have a distinct magnetic fabric (PORRECA et al. 2003, SAITO et al. 2003, TANAKA et al. 2004). Lahar deposits usually have no such magnetic fabric due to their low transportation and deposition temperatures (Table III, 6). Pyroclastic flow deposits have usually a loose texture, contain less mud in their matrix, and are less compact than lahar deposits. When a pyroclastic flow travels through an ice cap, it induces melt water and generates large volume lahars. In this type of scenario, the transition between pyroclastic flows and pumice-rich lahars ) can be gradual along the transporta- tion axis (CRANDELL 1987, BRANTLEY and WAITT 1988, SCOTT 1988, SIEBE et al. 1993). In the similar way, transition between volcanic debris avalanche deposits and clay-rich lahars is more or less gradual and complex (CALVARI et al. 1998, NEHLIG et al. 2001, CAPRA et al. 2002). Typical volcanic debris avalanche deposits are larger in volume, contain hum- mocks as the result of megaclasts transportation, and display a more chaotic internal architecture than lahar deposits. Till deposits may also look very similar to lahar deposits. However, tills generally do not incorporate vegetation fragments and do not expose lateral facies variations typical of lahar deposits. Finally, landslide deposits (that are sometimes tex- 164 Figure 7.5. Facies types of laharic deposits after VALLACE 2000: p. 611, fig. 7 turally similar to lahar deposits are generally very localised, and can be often associated to scars or other morphological features (headwalls, ridges, etc) pointing towards the source of the unconsolidated material. These small scale, landslide- associated diamictons can be easily distinguished in the field from lahar deposits by detailed surveying, careful mapping supported by aerial photo coverage if necessary. Lahars are among the most dangerous volcanic hazards. Lahars commonly reach unexpectedly inhabited areas destroying villages and cities. Ironically in many places worldwide lahar inundated areas re-inhabited, even larger scale than before the volcanic disaster stroke (Table IV, 1). Reduction of volcanic hazards by lahars commonly facilitated by huge effort to construct dams that may break the mechanic energy of the current, and hold back larger clast to travel long distances (Table IV, 2). Volcanic debris avalanches Volcanic debris avalanches are products of major sector collapses affecting a volcano under water (or fluids-)-satu- rated conditions (SIEBERT 1984, GLICKEN 1991, SMITH and LOWE 1991, UI et al. 2000). Debris avalanches are very rapid, inertial, granular flows, generally resulting from giant landslides occurring on an unstable portion of a mature volcanic edifice. Sector (or flank) collapses results in horse-shoe shaped amphitheatres from where unstable lava and brecciated pyroclastics are being removed and transported as a chaotic avalanche (Table IV, 3 and 4). Destabilization of the vol- cano’s superstructure by a shallow intrusion of magma high into the edifice is a process that has been linked to some of these collapses (e.g., Mount St. Helens in 1981) (Table IV, 5), while other debris avalanches have occurred without any production of juvenile magmatic materials (e.g., Bandai in 1888). Historic debris avalanches not linked to a magmatic eruption are commonly associated with enhanced fluid circulations and pressurization in shallow hydrothermal systems at the time of the collapse (REID 2004). Three main types of volcanic debris avalanches are recognized in the literature (UI et al. 2000), each of them refering to a different eruption style. Bezymianny-type debris avalanches refers to the blast- induced flank collapse of Mt Bezymianny in 1956 (Kamchatka, Russia; BELOUSOV 1996). This type of volcanic debris avalanches is triggered by a very powerful magmatic eruption. Bandai-type volcanic debris avalanches refer to the mass flow generated by a major phreatic eruption that decapitated Bandai volcano in Japan in 1888 (YAMAMOTO et al. 1999). In this later case, no juvenile material was found in the resulting volcanic diamicton. Finally, Unzen-type volcanic debris avalanches refer to the mechanical destabilization of an old parasitic dome situated on a steep flank of the erupting Mt Unzen in 1792 in Japan (UI et al. 2000). This type of catastrophic collapse is related to earthquakes rather than volcanic activity. Deposits from volcanic debris avalanches comprise two major facies types (UI et al. 2000): a block (or megaclasts- rich) facies, and a matrix facies (Figure 7.6). A debris avalanche block (or megaclast) is a fractured (and sometimes deformed), commonly overturned piece of the source volcano. Such blocks can reach tens to hundreds of metres in diameter, and may form an obscure stratigraphical section appear- ing in an exotic place (CACHO et al. 1994, Reubi and Hernandez 2000, Collot et al. 2001). Usually these blocks show many evidences of mech- anical stress, and internal fracturing can be very significant. Along the milled and strongly fragmented clasts, jig-saw fit textures are common (Figure 7.7). The matrix of debris ava- lanches is generally chaotic in texture (Table IV, 6), finely grained, and can host various lithologies representative of the sampled parts of the source vol- cano (REUBI and HERNANDEZ 2000). Areas affected by volcanic debris ava- lanches commonly show a typical hummocky morphology (Table V, 1 and 2) resulting from the distal trans- portation of fragmented megaclasts 165 Figure 7.6. Lithofacies types and changes recognized in the volcanic debris avalanche deposits of the Taranaki volcano, New Zealand (PALMER et al. 1991: p. 95, fig. 7). On "A" facies variations from source to distal areas demonstrated. On "B" a theoretical cross section of a volcanic debris avalanche deposit is shown(PONOMAREVA et al. 1998, CLA VERO et al. 2002, LE FRIANT et al. 2002). Hummocks commonly reach tens of metres across and no apparent distri- butional pattern has been identified yet (SIEBE et al. 1992). One of the main characteristic of volcanic debris avalanche deposits is their common fracture pattern. Jig-saw fit cracks are common on larger, brittle clasts such as volcanic lithics (e.g. fragments from coherent lava flows and domes). Jig-saw cracks of an avalanche block are joint-like and irregular, and usually remain closed. As megaclasts are transported over long dis- tances (i.e. >15 km), joints can open up and be filled progressively by finely grained matrix sediment coming from the milled part of the same fractured blocks (Figure 7.8). Despite the abundant geological evidence showing that volcanic debris avalanches are a natural consequence of volcano growth and destruction, their actual significance has just been recently emphasized. Volcanologists worldwide started to focus their attention on volcano col- lapses and the generation of debris avalanches after the catastrophic and well documented collapse of Mt St. Helens in 1980 (WAITT et al. 1983, DONNADIEU et al. 2001). Since then, volcanic debris avalanche deposits have been identified around many major volcanic complexes located in many different geological settings. The significance of volcanic debris avalanches in the evolution of a volcano therefore is increasingly recog- nized. Volcanic collapses generating debris avalanches leave very typical morphological features in the landscape. The most recognisable features are: (1) a large scar — or amphitheatre — located at the source of the ini- tial avalanche on the upper part of the volcanic edifice, often partially filled by subsequent lava domes, and (2) a widespread hummocky sur- face in the medial or distal reaches of the deposit. The hummocky region include blocks (megaclasts) of various sizes that are usually larger and more closely spaced in the proximal area of deposition. Recognition of hummocky surfaces however is not a sufficient criteria to confirm the volcanic debris avalanche origin of a diamicton in the field. Cross-sec- tional views of individual hummocks are usually required in order to identify fractured blocks, exotic volcanic stratigraphy and lithologies (Table V, 1). The recognition of a jig-saw fit texture affecting large coherent lava bodies embedded in altered matrix is necessary to infer the hummocky origin of small hills. Natural levees exceeding a few tens of metres in elevation are commonly reported in volcanic debris avalanche deposits such as the Socompa volcano in northern Chile (WADGE et al. 1995). Such cliffs are usually well developed in distal areas where the debris avalanche reaches flat lying areas of the surrounding ring plain (SIEBE et al. 1992). Slided and tilted blocks (or megaclasts) in debris avalanche deposits are more or less intact parts of the original volcanic edifice. Despite the fact that such blocks preserve the basic texture and lithology of the orig- inal portion of the volcanic cone affected by destabilization, they are often truncated and deformed, especially in distal regions, reflecting the impact of transportation over long distances. Very large blocks can significantly erode the substratum during the emplacement of the mass flow, and they are frequently mantled by fractured and crushed clast fragments where hummocks are formed. The magnetic fabric of a single avalanche block (megaclast) is usually fairly uniform. However, the magnetic fabric orientation can be very vari- able block by block, indicating uniform, en masse displacement of the source rock units into homogeneous megaclasts. The avalanche matrix is a mixture of finer grained sediment, derived from source rock lithologies. The magnetic fabric of the matrix is very randomly oriented, as a reflection of the well fragmented and diversified source rock. Debris ava- lanche matrix commonly contains colluvial and/or fluvial fragments picked up during travelling. Such fragments can be deformed and squeezed, forming flame-like structures injected between larger clasts. This is a common occurrence with water-saturated clasts of sedimentary rocks. The base of debris avalanche deposits is commonly fine grained, and has 166 Figure 7.7. Jig-saw fit texture of large volcanic lithic fragments of a volcanic debris avalanche deposit of the Ruapehu volcano, New Zealand Figure 7.8. Milled clast between large volcanic lithic clasts of a volcanicdebris avalanche deposit from the Ruapehu volcanomany textural features indicating high shear stress on the sole of the flow (Table V, 3). Soft sediment deformation, uni- directional flame structures, and other dewatering structures reflect loading and shearing in the base of the flow (Table V, 3). An increasing number of large volume volcanic debris avalanche deposits have been identified recently in many dif- ferent volcanic settings. The largest known volcanic terrestrial debris avalanche deposit has been identified around Mt Shasta in the Cascades (USA), where the deposit reaches at least a run-out distance of 45 km (UI and GLICKEN 1986). Volcanic debris avalanches have been extensively studied in volcanic arc settings, such as Unzen, Hokkaido-Komagatake (both in Japan), or at many Chilean or Mexican volcanoes. As previously indicated, volcanic debris avalanches can trans- form into clay-rich (“cohesive”) lahars, as it has been originally demonstrated for the Osceola Mudflow at Mount Rainer (US) (VALLANCE and SCOTT 1997). After travelling a distance of 2 km, the initial avalanche hit a major river valley and travelled an additional 120 km as a cohesive lahar (V ALLANCE and SCOTT 1997). Therefore, volcanic debris avalanches can affect very large areas, and can drastically modify the landscape. Despite numerous studies reporting the occurrence of volcanic debris avalanches deposits worldwide, only the Mt St Helens blast-triggered collapse was directly observed and well documented in recent historic times. The geological record of Quaternary volcanoes in Japan indicates that vol- canic debris avalanches occurred at least once in every hundred years (UI et al. 2000). In ancient terrains, identification of volcanic debris avalanches is more difficult to achieve. Volcanic debris avalanch- es can displace entire rock units from the source volcano as megaclasts measuring hundreds of metre across. Such very large blocks can be preserved and mimic original “in situ” volcanic successions. In this context, differentiation between “in situ” and allochton units is hindered by superimposed tectonic and metamorphic processes. The high mobility of these fluid-saturated mass flows may also result in the displacement of large sections of a cone tens of kilometres away from its source. In such a situation, stratigraphic and lithological correlations of avalanche block units with source rock units are difficult to establish. Detailed mapping is essential to delineate the avalanche boundaries, and their exact posi- tion in an exotic rock environment. Volcanic cone collapses and associated large volume debris avalanches have been well described in ancient volcanic settings such as the Cantal massif in central France (CANTAGREL 1995, REUBI and HERNANDEZ 2000, NEHLIG et al. 2001). Some large scale cone collapses are also related to major tectonic activity, such as movement of strike-slip faults beneath a volcano (LAGMAY et al. 2000). CECCHI et al. (2004) and ACOCELLA (2005) inferred that volcano spreading and associ- ated collapses (CECCHI et al. 2004, ACOCELLA 2005) can be enhanced by favourable substrate conditions beneath an over- steepened cone, usually dominated by the accumulation of lava domes. Many volcanoes in Northern Chile, where vol- canic edifices grown over salar deposits, illustrate this genetic relationship (Table V, 3, 4 and 5). Amphitheatres open at the head of volcanic valleys such as the Valle del Bove on the eastern, sea-facing flank of Mt Etna, are thought to have been shaped by catastrophic collapses (CALVARI et al. 1998, PARESCHI et al. 2006). Large scale avalanches are generated offshore by these giant landslides, as a result of unsupported parts of the growing volcanic edi- fice collapsing into the sea (CLOUARD et al. 2001, CLEMENT et al. 2003, MILIA et al. 2003, HURLIMANN et al. 2004, HILDENBRAND et al. 2006). Major volcano collapses have been similarly suggested to explain the formation of large half depression features found around volcanic islands such as Tenerife (Table V, 6) (HURLIMANN et al. 1999), and Gran Canaria in Spain (MEHL and SCHMINCKE 1999). However the general lack of associated volcanic debris avalanche deposits may complicate the interpretation of these depressions. Around volcanic islands however, bathymetric surveys often reveal the existence of an irregular seafloor morphology that could be interpreted as the result of submarine land- slides (KRASTEL and SCHMINCKE 2002, MASSON et al. 2002) and/or volcanic debris avalanches feeding submarine turbid- ity currents in their distal portion (SCHNEIDER et al. 2004). Such deposits have also been identified around most of the Hawaiian volcanic islands and around Reunion, in the Indian Ocean (OEHLER et al. 2004). The collapse of island vol- canic edifices usually leaves a subaerial to submarine horseshoe-shaped scar. On the island of Ischia for instance, the southern flank of the volcano was affected by a large scale collapse that generated a debris avalanche incorporating thou- sands of giant blocks, dispersed as far as 50 km offshore on the sea floor (CHIOCCI and DE ALTERIIS 2006). The transi- tion of volcanic debris avalanche deposits into submarine mass flow deposits has been well documented for many vol- caniclastic sequences, such as around Gran Canaria, Spain (SCHNEIDER et al. 2004). However, the exact interpretation of the succession of volcano-sedimentary events (e.g. primary debris avalanche deposits or secondary origin via reworking and redistribution) are in many cases not straight forward. In the Carpathian arc in Cenrtal Europe, volcanic debris avalanche deposits have recently been recognized. Large horseshoe-shaped morphological features are found in many Miocene to Pliocene erosional remnants of stratovolcanoes. Recent detailed studies suggest these residual morphologies are the result of major ancient volcanic collapses (KARÁTSON 1999, KARÁTSON et al. 1999). Since the initial recognition of the large collapsed crater zones, associated volcanic debris avalanche deposits have been identified. The authors suggest that an exhumed volcanic amphitheatre and the related vol- canic debris avalanche deposits may have been a cause for a major curvature on the Danube river channel (Table VI, 1), called Danube Bend (KARÁTSON et al. 2006). In addition, volcanic debris avalanche deposits in the Miocene Inner- Carpathian Volcanic Chain have been also identified in the Visegrád (Table VI, 2), Börzsöny and Mátra Mountains so 167far (KARÁTSON et al. 2001). Volcano collapse and associated volcanic debris avalanche deposits have been also identified recently in the Outer Carpathians and their potential role to host mineral deposits has been emphasized (LEXA et al. 1999, SZAKÁCS and SEGHEDI 2000, SZAKÁCS and KRÉZSEK 2006). Relationships between lahars and volcanic debris avalanches The idea that volcanic debris avalanches transform into lahars is now commonly accepted (LECOINTRE et al. 2002). When a volcanic edifice collapses, the resulting mass flow is often partially channelled by a major river system, and the avalanche becomes progressively more diluted downstream (CARRASCO-NUNEZ et al. 1993). This is especially the case for “dry” debris avalanches that have the capacity to incorporate large amounts of water during their emplacement in a river catchment. The flow can transform into various type of lahars s (PALMER and NEALL 1989, CARRASCO-NUNEZ et al. 1993, ENCINAS et al. 2006). Such mass flows usually contain megaclasts, representing rock types that are found in the upper volcanic edifice such as blocky fragments of lava domes. Deposits resulting from this type of lahar could be very similar to those left by lahars initiated by volcanic eruptions or meteoric events, especially in the distal zone of the ring plain. In many cases, volcanic debris avalanches and lahar currents are closely related. Debris avalanches form an end member in the spectrum of granular mass flows characterized by a very high sediment-to-water ratio in the current (Figure 7.3) (SMITH and LOWE 1991). Under-saturated debris avalanches are able to transform, with increasing water con- tent along a river channel, into debris flows, hyperconcentrated flows and finally, into dilute stream flows. In “dry” debris avalanches (i.e. clay-poor), the main support mechanism for clasts and particles is the grain dispersive force, which is gradually decreasing in strength once transformation into debris flows and hyperconcentrated flows occur. At the end of the spectrum, turbulence dominates in dilute stream flows (Figure 7.3). However, matrix cohesion of the initial mass flow may prevent this evolution. Grain support mechanisms in debris avalanches and debris flows are very strongly related, to the clay content of matrix in the current. Typically, debris avalanches sourced from hydrothermally altered portions of a volcanic cone are fluid-saturated (very high pore pressure), and will transform rapidly into “cohesive” debris flows that will not dilute further downstream (LECOINTRE et al. 2002). Fluid buoyancy also plays a significant role in particle sup- port for grain size particles, especially in hyperconcentrated flows. This process is less efficient in debris flows, and has now role in debris avalanches, as the relative coarse sediment to water ratio increases. The deposition mechanism of debris avalanches is predominantly en masse freezing, similar to cohesive debris flows. This differs notably from hyper- concentrated flow and dilute stream flow where grain-by-grain settling tend to dominate the deposition process. Lahars and fluvio-lacustrine depositional systems Lahars and catastrophic flood events can be triggered by major plinian eruptions (MANVILLE et al. 2005). Large vol- ume pumiceous deposits can block major water courses and may create temporary barriers such as tephra dams, result- ing into the development of extensive lakes (MACIAS et al. 2004). When the dams break, major volcaniclastic floods are initiated, and the generated lahars may cause significant changes on the landscape and devastation well after the volcanic eruption (MANVILLE et al. 1999, WAYTHOMAS 2001). Major pyroclastic events, such as the 12,900 yr BP Laacher Sea eruption in Germany, can modify the drainage system of an area, cause damming of major tributaries, and later trigger major floods once temporary water storages fail (SCHMINCKE et al. 1999). Similar reshaping of a large hydrographic net- work has been well documented for the 1.8 ka pumice-rich Taupo eruption (SEGSCHNEIDER et al. 2002a), where major re- sedimentation processes were particularly active in distant river catchments (SEGSCHNEIDER et al. 2002b). Lahars can leave along their route large volume of water-saturated sedimentary valley fills and overbank deposits. On the ring plain, such lahar deposits may create semi-permanent lakes (years to decades) that are gradually filled by fine grained sediment eroded from the barren surface. The gradual reestablishment of the original fluvial network progres- sively carve valleys into the lahar deposits, generating complex facies associations. In ancient volcanic settings, such flu- vio-volcaniclastic successions can be very difficult to decipher. On vertical sections, stratigraphic sequences showing sudden facies variations between debris flow, hyperconcentrated flow and normal stream flow deposits can be interca- lated with normal lacustrine beds (Figure 7.9). In horizontal view, lacustrine beds may be found as incorporated lenses into more massive debris flow diamictons-, alternating with sandy hyperconcentrated flow deposits. Because of intense erosion along the valley margins, diamictons resulting from late lahar activity and associated fluvial sequences may dis- play sharp and steep erosional contacts with earlier debris flow and hyperconcentrated flow deposits. Lahars can also be generated by relatively small-volume pyroclastic eruptions, such as phreatomagmatic explosions that trigger ground-hug- ging density currents. Such pyroclastic flows may enter a maar lake and transform rapidly into a mudflow, a process described for the Peperino Albano phreatomagmatic eruption from Colli Albani volcano, located 30 km to the south-east of Rome, Italy (GIORDANO et al. 2002). Large scale, basin-wide volcaniclastic successions associated with Large Igneous 168Provinces have been previously interpreted to be made of lahar deposits resulting from major destructional phases of quickly grown volcanoes such as the volcanic systems of Middle Jurassic Prebble and Mawson For- mations, Antarctica (ELLIOT 2000, ELLIOT and HANSON 2001). These deposits have been however recently re- evaluated, and re-interpreted as a direct result of phreatomagmatic explosive events, and being rather primary in their origin (WHITE and MCCLINTOCK 2001, ROSS et al. 2005, ROSS and WHITE 2005) Erosion of volcanic landforms Erosion of volcanic landforms has been studied with many different approaches. Sediment loss and sur- face modification that lead to the development of an erosional volcanic landforms can be a very efficient process. The quantification of the erosion rate is how- ever quite complex. Monogenetic volcanoes are an ap- propriate target when the objective is to quantify the erosion of small volume volcanic landforms. The process of erosion on scoria cones has been well docu- mented (Table VI, 3), and emphasizes a specific relationship between the height of the edifice, the opening of the crater, and slope angle changes (WOOD 1980b, 1980a). Such morphometric studies usually associate specific residual shapes of scoria cones (Figure 7.10) with a peculiar stage of evolution (PORTER 1972, WOOD 1980b, DOHRENWEND et al. 1986, HOOPER and SHERIDAN 1998). Various authors suggested that scoria cones can be almost totally flattened after a million year of erosion history, low in slope angle and their crater more or less unrecognizable (WOOD 1980b, HOOPER and SHERIDAN 1998, VESPERMANN and SCHMINCKE 2000). Such studies however have been derived from volcanic fields that evolved in arid climate, and where scoria cones are relatively simple in morphology, as well as uniformly composed of coarse-fine lapilli and ash beds (DOHRENWEND et al. 1986, SIEBE 1986, INBAR et al. 1994, HEIZLER et al. 1999, INBAR and RISSO 2001). More recent studies have demonstrated in addition that erosion stages for such volcanic landforms can be very much predetermined by their original morphology. In some cases, that original morphology may have been hindered by eruptive activity (e.g. phreatomagmatism) that led to the formation of a wide and steep crater in the cone (NÉMETH et al. 2005, MARTIN and NÉMETH 2006). The transition between cinder cone and a small volume composite volcano is very often gradual, and many erosional models cannot be directly applied in this situation (MCKNIGHT and WILLIAMS 1997). Also, many scoria cones go through an eruptive phase dominated by vigorous lava fountaining that produce thick weld- ed tephra units (NÉMETH 2004). Such units can also react to erosion as hard layers, and modify the expected erosional path of the corresponding volcanic landform (Table VI, 4). In this scenario, erosion may be driven by undercutting and small collapses on the outer crater rim, rather than chemical dissolution rain-generated debris flows, and/or eolian abra- sion (Table VI, 5) (NÉMETH et al. 2005). The situation is more complex for those small volume volcanoes, where the erup- tive history included a significant phreatomagmatic phase (NÉMETH 2004). Erosion of maars, tephra rings and tuff cones has not been studied in detail yet, and they seem to follow different rules to those applying to scoria cones. Tephra gen- erated by phreatomagmatic activity can form a “concrete-like” deposit due to a sudden loss of water, and behave as hard layers, resistant to erosion (WHITE 1991a). In such beds, erosion and especially the development of gullies (Table VI, 6) play an important role in the redistribution of volcanic material (WHITE 1991a). Phreatomagmatic activity produces also large volumes of water-saturated tephra. Those tephras can be easily remobilized during the eruption, forming major debris flows that carve a gully network on the flanks of the volcano (NÉMETH and CRONIN 2006). As previously men- tioned, debris flows tend to transform into hyperconcentrated flows in distal areas, where accumulations of tabular debris flows and hyperconcentrated flows units are intercalated with thin base surge and phreatomagmatic tephra beds (SOHN 1996, SOHN and PARK 2005, NÉMETH and CRONIN 2006). Thick accumulations of syn-volcanic, reworked tephra units can form a collar of flat lying volcaniclastic succession around a phreatomagmatic volcano (SOHN and CHOUGH 1989, SOHN 1996, SOHN et al. 2003, SOHN and PARK 2005). Due to the low slope angle and the broad opening of the vent, reworked volcaniclastics accumulate quickly in their crater (Table VI, 7) (WHITE 1991b). Under favourable conditions, crater lakes can be sedimentary traps of volcanic and non-volcanic debris. In extreme cases, maar-type volcanoes can have a crater hundreds of metres deep that can act as an active sedimentary basin over tens of thousands of years (WHITE 1992). During the initial erosional stage, entire blocks from the maar or tuff ring can collapse into the lake and initiate subaqueous 169 Figure 7.9. Diagram of syn and inter-eruption sedimentary facies may develop around a long lived strato-volcano after SMITH 1991: p. 113, fig. 3debris flows and turbidity currents (WHITE 1992, BÜCHEL and LORENZ 1993, FISHER et al. 2000). The tephra rim in many cases can be quickly eroded (thousands of years), and the maar lake acts as a sheltered, deep sediment trap (DROHMANN and NEGENDANK 1993, MINGRAM 1998, ZOLITSCHKA et al. 2000). Studies of lithofacies associations from eroded phreatomagmatic volcanic fields can help to reconstruct the erosion history of the landscape where the volcanoes erupt- ed (NÉMETH 2001). This long-term erosion rate studies suggest that lowering of the base level is in the order of a few tens of metres per million year for regions affected by a continental climatea result comparable with those previously pub- lished (NÉMETH and MARTIN 1999, NÉMETH 2001, 2003). However, (NÉMETH et al. 2006b) recently proposed that erosion 170 Figure 7.10. Stages of erosion (from "a" to "h") of scoria cones modelled by HOOPER and SHERIDAN (1998)rate calculations should be done with great care, and that a correct lithofacies study of erosion remnants should be com- pleted before attempting a detailed erosion calculation (Table VII, 1). In composite volcanoes, erosion is far more complex than in small volume volcanic edifices (WOOD 1978). Strato- volcanoes are generally long-lived, and extended periods of quiescence separate eruptive phases (DA VIDSON and DE SILVA 2000). Periods of eruptive activity can also be very complex in terms of eruption style and intensity. Freshly produced tephra can be quickly remobilized into syn-eruptive lahars, commonly accompanied by primary pyroclastic density cur- rents (DA VIDSON and DE SILVA 2000). In inter-eruptive periods, and prior to any recolonization of the bare landscape by vegetation, rain-induced lahars can redistribute large volumes of loose volcanic sediment onto the ring plain. As a result, especially on stratovolcanoes with low frequency eruptions and long periods of quiescence, the continuously eroded edi- fice is progressively surrounded by a growing and thickening succession of volcaniclastic sediments (LECOINTRE et al. 1998, DA VIDSON and DE SILVA 2000, LECOINTRE et al. 2004). Eruptions, coupled with intermittent erosional phases on a volcano that is active over hundreds of thousands of years, are able to generate a ring plain hundreds of metres thick, and mostly dominated by reworked volcaniclastic sediments (Table VII, 2). Numerous long-lived composite volcanoes are surrounded by such a volcanogenic sedimentary pile. The detailed study of the corresponding stratigraphies help to understand the nature of the depositional processes for each volcanic succession found in the ring plain, and may pro- vide vital information about early evolution stages of the central cone. Very often, products on the cone have been already removed by erosion, and only preserved by redeposition in the lowlands of the ring plain. Erosion of composite cones has been studied by applying a range of complementary methods on many different volcanic settings. From the Carpathians, Miocene to Pliocene stratovolcanoes (Table VII, 3) have been selected to quantify the main factors control- ling landform modification, and identify long term erosion rates (KARÁTSON 1996, 1999). Erosion of these volcanoes resulted in a significant widening of their central craters, forming a talus bordering depressions similar to calderas (termed erosional calderas). Erosion of composite volcanoes can generate very complex structures, especially when flank or sector collapse on the volcano occurs in its history (Table VII, 4). THOURET (1999) demonstrated that on many lava dome-dominated arc volcanoes, the landscape evolution is predominantly controlled by pulsating growth and successive collapses (THOURET 1999). Such volcanoes can develop largely as an asymmetric cone with truncated flanks (Table VII, 5). In this context, flank and ring plain sections show widespread hummocky surfaces determined by the distribution of major volcanic debris avalanches, such as the Calbuco volcano in southern Chile. Erosion of volcanic landforms can produce spectacular micro, meso- and macro-features. 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W. 2000: Eifel maar lake sediments: Calendar-year archives of past environments. — Terra Nostra 6, pp. 545–549. 177178 Chapter 7 Location map 1 — Shiveluch, Russia 2 — Ruapehu, New Zealand 3 — Casita, Nicaragua 4 — Pinatubo, Phillippines 5 — Ambae, Vanuatu 6 — Nevadodel Ruiz, Columbia 7 — Pico de Orizaba 8 — Mt Rainer, Washington, USA 9 — Cotopaxi, Ecuador 10 — Villarica, Chile 11 — Bandai, Japan 12 — Mt St Helens, Washington, USA 13 — Bezymianny, Russia 14 — Unzen, Japan 15 — Taranaki, New Zealand 16 — Tongario, New Zealand 17 — Calbuco, Chile 18 — Mt Shasta, California 19 — Hokkaido-Komagatake, Japan 20 — Cantal, France 21 — Valle del Bove, Etna, Italy 22 — Tenerife, Spain 23 — Gran Canaria, Spain 24 — Hawaii, USA 25 — Renuion, France 26 — Danube Bend, Hungary 27 — Börzsöny Mts, Hungary 28 — Mátra Mts, Hungary 29 — Pali Aike Argentina 30 — Al Haruj, Libya 31 — Hopi Butte, Arizona, USA 32 — Cerro Colorado, Sonora, Mexico 33 — Laacher See, Germany 34 — Lake Taupo, New Zealand 35 — Colli Albani, Italy179 Chapter 7 Plate I 1. A) Volcanic ring plain around Ruapehu volcano in the central North Island, New Zealand. B) Volcanic ring plain of the Karymsky volcano (Russia) evolving in a caldera. 2. Tephra succession around the Tongariro–Ruapehu volcanoes, New Zealand. 3. Lahar deposits in the Ruapehu ring plain. The deposit consists of large rounded blocks of volcanic lithics in fine-grained matrix. The deposit is located about 15 km from the source. 4. Columnar jointed block in block and ash flow triggered lahar deposit in Ruapehu. 5. Snow capped peak of Ruapehu a perfect site to initiate lahars in case of hot pyroclastic flows would travel through this region. A combination with an existing crater lake in the peak of Ruapehu makes this volcano hazardous of lahar initiations. 180 Chapter 7 Plate II 1. Deep valley network filled with ancient lahar deposits in the western flank of Ruapehu. 2. Block and ash fan on the Calbuco volcano in southern Chile dominated by block and ash flow deposits and minor reworked volcaniclastic successions. 3. Debris flow deposit in the Osorno volcano ring plain in southern Chile. 4. Hyperconcentrated mud flow deposits from the 18th of March 2007 lahar from Ruapehu volcano, New Zealand, near the Tangiwai Bridge. 5. Lahar destroyed Aztec city of Cholula in the ring plain of Popocateple in Mexico. Popocateple is visible in the distance. 6. Aztec pyramids under matrix supported distal lahar deposit facies in Cholula, Mexico.181 Chapter 7 Plate III 1. Non-charcoaled bed flattened tree trunks in the1971 Turbio Valley lahar deposits near Pucon, Chile. 2. Massive, thick unit of volcaniclastic breccia in the Liucura Valley near Villarica volcano in Chile. The origin of the sequence is still under debate. The textural characteristics is similar to debris flow deposits from lahars, however, the angular and largely monomict volcanic lapilli content suggestive for pyroclastic flow origin. It is a good compromise to interpret this section as a pyroclastic flow triggered lahar deposit. 3. Close up view of a hot lahar deposit (same as in Plate III, 2). 4. Fluvial deposits of the Turbio valley. These valley is a common routes of lahars such as the 1971 lahar initiated from the NE flank of the Villarica volcano, Chile. 5. Turbio valley filled by the 1971 lahar deposits that formed a delta in the Lake Villarica, Chile. 6. Block-and-ash flow deposit texturally similar to lahar deposited debris flow deposit, but have characteristic magnetic fabric as AMS study reviled its magnetic fabric (Mendoza, Argentina).182 Chapter 7 Plate IV 1. Ring plain around Popocateple volcano in Puebla, Mexico. In the historic time devastating lahars inundated large Aztec cities in the region, regardless of the complete destruction of those ancient cities, Cholua and other large cities, including Mexico City of 15 million people, are in potential danger of lahars may initiate from Popocateple. 2. Sebu dams built in many Japanese volcanoes, to reduce the mechanical energy of down flow lahars. 3. Horse-shoe shape escarpment of the Mt St Helens volcano on an oblique satellite image (Google Earth) as a result of sec- tor collapse of the summit of the volcano. Note the flat smooth surface area in the ring plain partially filling a lake (black field). 4. Horse shoe-shaped escarpment of the summit of Mt St Helens, as a result of the 1980 eruption (photo by S. J. Cronin). 5. Mt St Helens collapsed summit area looking from the ring plain (photo by S. J. Cronin). 6. Debris avalanche matrix with chaotic volcanic lithic frag- ments from a volcanic debris avalanche deposit of the Ruapehu volcano, New Zealand.183 Chapter 7 Plate V 1. Proximal part of a volcanic debris avalanche with large (tens of metres across) hummocks about 15 km from the crater of Popocateple volcano. 2. Distal part (about 50 km away from the source) of hummocky surface with small (metre scale) hummocks around Calbuco volcano, southern Chile, as a result of a major volcanic debris avalanche of this volcano. 3. Sheared base zone of a debris avalanche deposit on Popocateple, Mexico. 4. Large, oversteepened strato cones develop over salar deposits in the Altiplano, ready to collapse due to the lubricated basal rock units. 5. Hummocky surface of a volcanic debris avalanche deposit in the Altiplano developed over salar deposits. 6. Steep northern flank of Tenerife Island inferred to collapsed several times in its history. The present day edifice of Pico del Teide is over 4000 m above sea level and unsupported in the northern side of the island.184 Chapter 7 Plate VI 1. Danube bend is inferred to be a former horse-shape amphitheatre structure, a result of a former volcano collapse. 2. Volcanic debris avalanche deposits in the Visegrád Mt, Hungary. 3. Eroded scoria cone from Mendoza, Argentina. 4. Skeleton shape volcanic cones in the Pali Aike Volcanic Field in Argentina. The erosion dominated buy high wind, removed the loose tephra, leaving behind skeleton-like struc- tures dominated by lava spatter layers. 5. Undercutting of scoria cone beds causing sudden collaps- es of large part of the pyroclastic units of scoria cones of the Al Haruj al Abiyad Volcanic Field in Libya. This way, the cones erode in a different path than it has been predicted from scoria cones dominated by granular lapilli successions. 6. Deep gully network on the outer flank of the Cerro Colorado tuff cone in the Pinacate Volcanic Field, Sonora, Mexico. 7. Volcaniclastic sediment accumulated in a crater lake of a maar of the Hopi Butte, Arizona. Note the diverse texture of individual clasts and their rounded shape. 185 Chapter 7 Plate VII 1. General model after NÉMETH et al. (2006) of the erosion of maar volcanic complexes in western Hungary. (A) Subaqueous vent(s) in water filled maars. (B) Scoria cone and associated lava flow(s) erupted in a water filled maar. (C) Scoria cone erupted in a dry maar and fed lava lake. Continuous black lines represent potential erosion level in each case. Note the significantly different meanings of the preserved palaeo- surfaces covered by lava units. 2. Thick volcaniclastic succession in an incised valley 40 km from the Ruapehu volcano, New Zealand. 3. Eroded Miocene volcanic landforms of the Visegrád–Börzsöny Mtn. 4. Collapse scars (horse-shoe-shaped craters) of a lava dome domi- nated strato volcano in the Altiplano, Chile. 5. Irregular cone morphology of Calbuco volcano in southern Chile is a result of pulsating lava dome growth and collapse. This volcanic evolution and the erosion together produce a truncated cone mor- phology. 6. Micro-yardangs on vesicular basaltic lava flow surface of a pahoe- hoe lava field of Al Haruj al Abuyad, in Central Libya.186 Chapter 7 Plate VIII 1. Elongated mega-yardang developed on an ignimbrite surface in the high wind area behind the Andean mountain chain in Mendoza. 2. Close up view of a mega-yardang surface developed over an ignimbrite surface. Note the micro-yardang texture similar to those developed on vesicular basaltic lava flows in Libya (see Plate VII, 4). 3. Thermal wheatering forming polygonal fractures over basaltic lava flow surface in Central Libya. 4. Displaced fratured blocks of a lava surface due to thermal wheatering of basaltic lava flow in Central Libya. 5. Radial jointing pattern over large coherent lava body due to thermal wheatering of lava flow surface in Central Libya. 6. Large-scale redistribution of pumiceous fall deposits (white zones in the peak) from volcanoes in the high Andes forming dunes in the lee-side of the mountain ranges. Introduction Subaqueous volcanic eruptions are the most common on earth, considering the total volume of magma erupting along mid-oceanic rift axes, hot spot-related volcanoes and volcanoes erupting in shallow marine, lacustrine or subglacial envi- ronments. The total volume of effusive and explosive eruption products and their economic significance is huge in com- parison to subaerially erupted volcanic products. In spite of the abundance of subqueous volcanism on Earth, the under- standing of volcanic processes in such environments is very limited for two reasons; 1) logistical problems associated with collecting samples and studying sites underwater and 2) the long distances to and remoteness of the most suitable study sites. From a logistical point of view our access to subqueous eruption sites is very limited, and even in cases where good access is possible we are largely limited to observations from the water surface (e.g. direct observations of the eruption clouds that may breach the water surface). Even in the case of submersible techniques, visibility in water is much less then in air and detailed observations are very difficult to obtain. Also, sample collection from the seafloor (e.g. after eruptive events) is largely limited to dredging or coring, which do not allow identification of outcrop (tens of metres scale) observations of key sedimentary features that may be key to interpreting the transportation and depositional regime that formed the volcanic unit. The remoteness of many subaqueous eruptions also limits our ability to reach eruption sites quickly enough to see the main eruptive phases, commonly responsible for the formation of the majori- ty of the volcanic sediments and effusive products. Therefore our understanding of subaqueous volcanism relies main- ly on studies of ancient volcanic successions, where we are confidently able to establish the subaqueous environment where the studied volcanic rock unit formed (e.g. identification of subaqueous deposition for under- and overlying sed- imentary successions sandwiching volcanic rock units (WHITE et al. 2003). However, although this approach is widely used, potentially serious errors resulting from simple sedimentological problems that arise in the interpretation of con- tact zones between non-volcanic, marine and volcanic rock successions remain a source of debate in many studies. This is partially the reason why many Precambrian volcanic successions are used for understanding subaqueous volcanic processes, since at that time, the Earth’s surface was covered by water (MUELLER 1991, DOUCET et al. 1994, LAFRANCE et al. 2000, MUELLER et al. 2000, 2002). In spite of every effort to use careful, well-supported arguments and reason- ing to pursue such studies on older rock formations, we all face major difficulties in understanding the tectonic and palaeoenvironmental development of subaqueous volcanic settings. It is also common to find that although older vol- canic rock units may be well-suited to understanding subaqueous volcanism, they are commonly discontinuous or dis- sected, hindering any chance to see the full 3D facies relationships of different volcanic rock units (WHITE et al. 2003). This problem is one of the major issues that makes it difficult to confidently use older successions for facies modelling of subaqueous volcanic successions, even if they are relatively young (a few millions of years old). Effusive subaqueous volcanism Effusive volcanism in subaqueous settings is the most widespread type of volcanism on Earth, especially along the mid-ocean ridges (BATIZA and WHITE 2000). Along with extensive lava flows, large volumes of non-explosive, quench fragmented rock called hyaloclastite (Plate I, 1) forms due to non-explosive fragmentation of erupting lava (LONSDALE and BATIZA 1980, BATIZA et al. 1984, 1989, SMITH and BATIZA 1989, BATIZA and WHITE 2000). In addition, significant 187 Chapter 8volumes of magma is emplaced subaqueously as intrusive bodies (Plate I, 2) and associated peperite, formed by in situ mingling and mixing of fragmented shallow dyke and sill magma with host marine and/or lacustrine sediments (EINSELE 1986, WHITE et al. 2000, SKILLING et al. 2002, WHITE and HOUGHTON 2006). Magma rise near the sea/lake floor is predominantly controlled by the marine and lacustrine sediment pile. If the loose marine and lacustrine sediment pile is thick, magmatic crack propagation stops functioning and magma can stop rising and spreads laterally over the sea/lake floor, or can invade the marine/lacustrine sediment, forming extensive peprite facies (WHITE and BUSBY-SPERA 1987, MCPHIE 1993). Effusive and intrusive-dominated subaqueous volcanism produces large volumes of sediment, fragmented during non-explosive or weakly explosive events, that are distributed by currents or gravitational fail- ure events and commonly form sheet-like deposits on the sea/lake floor, inter-bedded with non-vol- canic marine and lacustrine sediments (WHITE and BUSBY-SPERA 1987). Seamounts are a common volcano type associ- ated with mantle up-welling sites, where magma pours out onto the seafloor, and forms submarine volcanic edifices (STAUDIGEL and SCHMINCKE 1984, MAICHER 1999, CORCORAN 2000, TRUA et al. 2002, CLOUARD et al. 2003, TARDUNO et al. 2003, KOPP et al. 2004). Large volcanic plateaus (e.g. Ontong- Java plateau) are also commonly associated with hot spot magmatism, and may form sheet-like flood lava fields and associated hyaloclastite units that cover thousands of square kilometres (TARDUNO et al. 1991, BERGER et al. 1992, KERR et al. 2000, WESSEL and KROENKE 2000, MANN and TAIRA 2004). The most obvious subaqueous lava forms are pillow lavas (Plate I, 3), lobate lava flows and sheet-like lava fields (BATIZA and WHITE 2000). Laboratory experiments show that at low effusion rates and on gentle slopes pillow lavas form (MOORE 1975, CAS 1992, BATIZA and WHITE 2000). Increasing effusion rates and steeper slope angles result in lobate lava flows and sheet-like lava fields (BATIZA and WHITE 2000). Pillow lavas have many distinct morphological features (Figure 8.1). Pillows can be elongate, oval shaped to complex interconnected piles and are commonly associated with hyaloclastite facies (WALKER 1992). Mega-pillows (Figure 8.2) have also been recognised (GOTO and MCPHIE 2004). The rim of the individual pillows is glassy, sometimes with multiple glassy layers (KAWACHI and PRINGLE 1988); their interior, however, is usual- ly more crystalline due to their slower cooling rate. Pillow lava piles commonly form complex pillow volcanoes. Subaqueous lobate lavas form large lobate-shaped, often sack-like forms that are interconnected. Sheet lava flows are associated with fast spreading mid-ocean ridges and are believed to fed by lava tube networks. Hyaloclastite is a glass-fragment rock that results from non-explosive, quench fragmentation of subaqueous lava flows (SILVESTRI 1963, BATIZA et al. 1984). Basaltic hyalo- 188 Figure 8.1. Theoretical architecture of a pillow lobe (after YAMAGISHI 1985 in GOTO and MCPHIE 2004: p. 322, fig. 13). Multiple crusts form on the tip of the advancing lava toe. Two individual lava lobe initiates from a single master lobe. The open and closed arrows refer to spreading and flow direction respectively. A) architecture of a pillow lava lobe, B) architecture of a subaquaeous lava lobe Figure 8.2. Diagrammatic representation of megapillow (A) and lobate lava flow (B)(after GOTO and MCPHIE 2004: p. 324, fig. 14)clastite is commonly associated with extensive pillow, lobate and sheet-like lava flows (BATIZA et al. 1984). In modern sea floors acidic hyaloclastite piles are relatively rarely identified and are mostly associated with rift set- tings. However, rhyolitic hyaloclastite is relatively com- mon in the geological record, mostly inter-bedded in marine sediments and associated with lava domes or cryptodomes that grew within and on top of sediments on the former seafloor (Plate I, 4) (CAS et al. 1990, YAMAGISHI 1991, SCUTTER et al. 1998, DERITA et al. 2001, RINALDI and VENUTI 2003, NÉMETH et al. 2007). These hyaloclastite-dominated rhyolites, rhyodacites and dacite domes can be very complex volcanoes (Figure 8.3) (MCPHIE and ALLEN 1992, MCPHIE et al. 1993, GIMENO 1994, GOTO and MCPHIE 1998, DOYLE and MCPHIE 2000, GIFKINS et al. 2002, RINALDI and VENUTI 2003). Subaqueous volcanic successions and associated rhyolite lava domes are common from a range of tectonic settings (CAS et al. 1990, GIMENO 1994, DERITA et al. 2001). Lava domes, cryptodomes and lava flows (Figure 8.4) easily burrow into and apparently rapidly expand into soft sedi- ments in subaqueous settings (KANO 1989, HANSON and HARGROVE 1999, GIFKINS et al. 2002), forming various types of peperite and associated hyaloclastite units. The term cryptodome is used to describe coherent bulbous magmatic bodies emplaced at shallow levels into host sediment that never breached the sedimentary cover (GOTO and MCPHIE 1998, STEWART and MCPHIE 2003). Cryptodomes in subaqueous settings are surrounded by intrusive hyaloclastites (STEWART and MCPHIE 2003) formed by the quench fragmentation of the margin of the intruding magmatic body (Figure 8.5 and 8.6). Intrusive hyaloclastite is also known as peperitic hyaloclastite (MCPHIE et al. 1993). Hence, intrusive hyaloclastite and peperite are very closely related terms. Some intrusive hyaloclastite bodies could also be termed peperite, but only if there is evidence of host sediments injected into the intrusive body. By contrast, while endogenous lava domes are partial- ly grown in soft sediment (MCPHIE et al. 1993, GOTO and TSUCHIYA 2004), at their upper margins lava breaches the 189 Figure 8.3. Theoretical model of the evolution of the Pálháza cryptodome- and dome-complex (after NÉMETH et al. 2007). In the initial stage rhyolitic magma intrude into the seafloor sedi- ment, forming intrusive hyaloclastite and various types of peperites (A). As new rhyolite magma intrudes, the new melt breaches the seafloor sediments, and pour out to the seafloor. Upon contact with the water, thick pile of hyaloclastite forms around the lava dome. Currents, and gradual increase of the slope angle of the growing dome complex initiates remobilisation, and reworking of the accumulated hyaloclastites Figure 8.4. Diagrammatic representation of facies relationships of subaqueous lava flows and lava domes (after MCPHIE et al. 1993). Numbers represent theoretical sections across the lava flows and lava domes. The sections are very similar, and made a challenging work to distinguish subaqueous lava flows from lava domes preserved in the geological recordsedimentary cover and, upon contact with the sea/lake water, quench fragmentation of the lava surface forms in situ hyalo- clastite. These hyaloclastite successions may laterally intercon- nect with redeposited hyaloclastite units formed by volcani- clastic deposits carried away from the in situ hyaloclastite piles by currents (GOTO and TSUCHIYA 2004). Extrusions in very shallow subaqueous environments may generate explosive eruptions and form tephra mounds and cones that overlie domes (CAS et al. 1990, 1992). Distinguishing geologic expo- sures of subaqueous cryptodomes from endogeneous lava domes may be challenging due to the fact that a cryptodome may expand enough to breach the host sediment. Cryptodome and dome structures can overlap with a highly intricate archi- tecture that can contain a variety of rock textures, as is documented from the Tokaj Mountains, NE Hungary (NÉMETH et al. 2007). Eruption dynamics of explosive subaqueous volcanism The basic difference between subaqueous and subaerial volcanism is that the magma, upon reaching the lake or sea floor, can encounter water (WHITE et al. 2003). Therefore, there are two basic problems we have to face (WHITE et al. 2003); 1) how this magma–water interaction took place, and 2) when, or what stage of the eruption was affected (most)? It is known from direct observations and inferred from ancient settings that many subaqueous volcanic eruptions occurred in shallow subaqueous environments where the eruption intensity was large enough to breach the water surface and produce eruption clouds that emerged above the water surface, or created physical conditions underwater that are similar to subaerial eruption clouds (WHITE et al. 2003). In summary, we can state that water differs significantly from air as an eruption media (WHITE et al. 2003) with respect to the following factors: 1) the role of steam generated during the eruption, 2) the role of pressure caused by the water column, 3) the role of heat capacity and conductivity of water, and 4) the water rheology. These four fundamental parameters are significantly different from eruption conditions in air. When magma rises and erupts in dry environments, its explosive fragmentation is driven by the exsolution and expansion of gases that were trapped within it under high pressures at depth. However, in wet environments, magma is dominantly fragmented through a conversion of thermal energy to mechanical energy when the >1000 o C magma meets water. This contact leads to a chain-reaction process often referred to as molten-fuel-coolant-interaction (MFCI) (ZIMANOWSKI et al. 1991, 1997a, ZIMANOWSKI 1998). The MFCI process not only pulverises and chills of the magma, but the resulting shock waves also permeate and disrupt surrounding rock and sediment (WOHLETZ 1986, ZIMANOWSKI 1998). Water (the coolant) vaporizes above its boiling point upon contact with hot magma (the molten fuel) and its expansion enhances magma fragmentation. Entrapped water can expand quickly and fragment the melt. During the pre- mixing stage between magma and water, the water becomes superheated. Thermohydraulic fragmentation of the melt 190 Figure 8.5. Theoretical facies distribution associated with a subaqueous cryptodome (after MCPHIE et al. 1993) Figure 8.6. Facies association of a subaqueously emplaced rhyolite lava flow from the Miocene, Ushikiri Formation, Japan (after KANO et al. 1991)take place and the fragments dispersed in the superheated water, followed by sudden fragment dispersion when the superheated water flashes to steam (ZIMANOWSKI et al. 1997a, b). The resulting steam affects the dispersion of the frag- mented magma during the lifetime of the steam (which is usually short due to its buoyancy and low density). Steam also has an insulating effect, which could help to form spatter and locally weld ejecta close to eruption sites, where no direct contact between hot particles and cold water occurs (MUELLER and WHITE 1992, KANO et al. 1994, MUELLER et al. 2000, CAS et al. 2003). In shallow water, water can boil to steam and expand dramatically. However, it is inferred that the expansion of and the phase change from water to steam is gradually suppressed with increasing water depth, and generally believed that at over 3 km water depth no meaningful phase changes are possible and only limited expan- sion can take place. The water pressure also increases with depth, which affects the solubility, expansion, and release of magmatic volatiles, as well as the development of steam (FISHER 1984, FISHER and SCHMINCKE 1984). In practical terms, this means that increasing pressure decreases magma fragmentation since volatile expansion is suppressed. In a simple way this means that much higher concentrations of magmatic volatiles are required in order to reach the same amount of fragmentation at greater depth (e.g. volatile-rich magmas) (DIMROTH and YAMAGISHI 1987). The role of pressure on magma–water interaction issues, however, remain controversial. Experimental studies demonstrate that at fixed magma–water ratios the violence of magma–water interaction increases with pressure (WOHLETZ 2003) (e.g. deeper water would cause stronger, more powerful MFCI interactions); however, this might be just a reflection of the experi- mental set up. Water heat capacity and conductivity is significantly higher in water than in air and therefore rapid cooling of magma and erupted gases are expected. The erupting melt will quench fragments quickly when direct contact between magma and water occurs. Further quench granulation can also take place as large hot particles break apart, leading to formation of smaller and smaller particles as the process continues until the magma has cooled completely. Intrusion into seafloor water-saturated sediments can result in quench fragmentation and production of large volume of spalled glass from the surface of the erupting magma. Already fragmented hot pyroclasts leaving a subaqueous vent can suffer further fragmen- tation in the margin of the eruption column, where individual clasts come into contact with ambient water. In general fluidal clasts (e.g. spindle bombs) are not expected to form in subaqueous environments due to the brittle fragmentation of the rapidly chilled fragments, unless water is temporarily excluded from eruption sites (MUELLER and WHITE 1992). Such situations are expected from time to time, when magmatic volatiles are able to exclude water temporarily or when significant volumes of steam are produced during the contact between the hot magmatic bodies and water. Still, if flu- idal particles are hot enough when they come into contact with ambient water, they can brittly fragment and generate blocks and shards from the original fluidal clasts (MUELLER and WHITE 1992, CLAGUE et al. 2003, MUELLER 2003). On the other hand, large-volume fluidal clasts can generate steam envelopes that insulate the clast from surrounding water and thus prevent further brittle fragmentation In comparison to air, water is dense and viscous, which strongly affects the eruption plume geometry as well as the dispersal of the ejected material (WHITE et al. 2003). Ballistically transported fragments are not expected to be common due to the high viscosity and density of the water in comparison to air (e.g. the erupted fragments will slow down, stall and fall back through vertical movement, instead of following a ballistic trajectory). Ballistic transport is only likely in near-vent water exclusion zones, or when the eruption takes place in shallow water, and fragments can breach the water surface (KOKELAAR 1986, WHITE 1996). Buoyant particles can be transported to the water surface, and then aqueous cur- rents can carry them away. Even basaltic (high density clasts) can be buoyant, at least while they are hot and have entrapped magmatic volatiles in their core. Subaqueously erupted pumice can float until their pores become water satu- rated and they sink (MANVILLE et al. 1998). Ash size particles can accumulate “en masse” from hot/warm water plumes as they drift away from the eruption sites (CASHMAN and FISKE 1991, FISKE et al. 1998). The resulting deposits of such aqueous plumes perhaps form different bedding characteristics than deposits settled from subaerial eruption plumes. It has recently been recognized that large volumes of fine particles from subaqueous eruptions, initially transported near the water surface, can form vertical density currents (MANVILLE and WILSON 2004). Once they reach the sea/lake floor these vertical density currents form radially dispersing, now horizontal, density currents (WHITE 1996, 2000). Perhaps in the case of drifted subaqueous ash plumes vertical density currents can form radiating density currents on the sea/lake floor, complicating the reconstruction of their source. Gas-supported subaqueous density currents are considered to be true subaqueous pyroclastic flows (WHITE 2000). To generate such currents, the system must have high particle concen- trations and have very low concentrations of low density, buoyant (e.g. pumice) particles in order to keep the current moving on the sea/lake floor. The higher density of the water as the supporting media will make these flows slow-mov- ing phenomena. The few field studies and theoretical considerations available suggest that the resulting subaqueous pyro- clastic flow deposits are probably very similar to subaerial pyroclastic flow deposits, even including significant welded successions (SPARKS et al. 1980). Subaqueous surges (e.g. turbulent, low particle concentration gaseous density currents), however, are not expected to move along the sea/lake floor due to their low density and therefore high buoyancy (WHITE et al. 2003). 191Shallow subaqueous explosive volcanism In shallow subaqueous environments where the water depth is less than few hundred metres, Surtseyan-type (predom- inantly basaltic) explosive subaqueous eruptions can take place (WHITE and HOUGHTON 2000). Surtseyan volcanoes are known from all shallow subaqueous environments (submarine, lacustrine, glacial). The low confining pressure of the water column contributes to the explosiveness of this style of volcanism, and is inferred to occur at water depths of up to about 200 m (WHITE and HOUGHTON 2000). Surtseyan-type volcanoes are generally monogenetic (see chapter 5) and their eruption duration ranges from few days to years (Plate I, 5) (WHITE and HOUGHTON 2000). Surtseyan-type eruptions share many characteristics with subaerial maar and tuff ring forming eruptions, particular- ly magma fragmentation as a consequence of magma–water interaction driven by MFCI processes. Magma–water inter- action in Surtseyan eruptions mostly takes place at the interface between the top of the conduit and sea/lake floor (KOKELAAR 1983, 1986, WHITE and HOUGHTON 2000). This means that during magma–water interaction abundant water and/or water saturated slurry is available to drive explosions (KOKELAAR 1983). Surtseyan eruptions build up steep sided tuff cones of phreatomagmatic tephra constructed by fallout and pyroclastic density currents (WHITE and HOUGHTON 2000). The volcanic edifices are surrounded by thick gravity mass flow deposits resulting from mass wasting of the loose and wet tephra (VERWOERD and CHEVALLIER 1987, SOHN 1995, COLE et al. 2001, MARTIN 2002, MAICHER 2003). Volcano collapses are also common and are recorded by large scars draped by younger tephras. The initial eruption of such volcanoes commonly starts with lava effusion, form- ing thick piles of pillow lavas (STAUDIGEL and SCHMINCKE 1984, MCPHIE 1995, ANDREWS 2003). Over these pillow volcanoes a partially subaqueous and partially subaerial tuff cone can develop (STAUDIGEL and SCHMINCKE 1984, CAS et al. 1989, SCHMIDT and SCHMINCKE 2002). Tuff cones that form over a pile of hyaloclastite and/or pillow mounds in subglacial settings are commonly named tindars (SMELLIE 2000). Magma fragmentation in shallow subaqueous settings is inferred to be similar to subaerial volcanoes, with decompression triggered by magma rise toward the surface (in this case, the sea- or lake-floor) where volatiles exsolve and form a magma-bubble foam in the upper section of the conduit. The general lack of accidental lithic fragments derived from the conduit wall indicates that fragmentation mostly takes place in the topmost part of the conduit or in the crater. Fragmentation of magma in different stages of vesiculation produces pyroclasts with wide-ranging but generally low vesicularity; this range of vesicularity is characteristic of Surtseyan deposits (HOUGHTON and WILSON 1989). The magma foam or pyroclast-charged slur- ry will exit the vent if its pressure exceeds the ambient pressure which is controlled by 1) the depth of fragmenta- tion relative to the sea/lake floor (WILSON and HEAD 1981) and 2) the volcanic conduit radius (WOODS 1998) (narrow- er conduits in hard rocks produce higher pressure). In this respect, in subaqueous environments it is inferred that the fragmented slurry-like material leaves the vent slowly and looks rather like a “boil over” event, especially if the erup- tion is characterised by low mass flux or occurs in deeper water (WOODS 1998). Direct observations confirm that dur- ing Surtseyan eruptions a gas + water + ash mixture radial- ly leaves the vent accompanied by a hemispherical shock wave manifested as an expanding spall dome (Figure 8.7) (WHITE 1996). During the onset of a subaqueous eruption a gas-thrust zone of the eruption plume forms over the vent, but this gas-thust region quickly transforms to a buoyant convective eruption column. This buoyant eruption column can reach the water–air interface, but is usually not able to breach it, unless the eruption is initiated in a very shallow subaqueous environment (KOKELAAR 1983, KOKELAAR and DURANT 1983). In deep water eruptions, an extensive mush- 192 Figure 8.7. A model for spall dome formation (in time sequence of “a” to “d”) during subaqueous explosive eruption, based on the reconstruction of the formation of the Pahvant Butte sublacustrine volcano (after WHITE 1996: p. 259, fig. 9a–d) room-like cloud may form near the water surface as a result of the rise of the buoyant convective cloud (CASHMAN and FISKE 1991). These near surface regions can quickly become overcharged in ash particles, and the particle cloud can fall back to the sea floor via vertically descending sediment gravity currents (CAREY 1997). Once these currents hit the sea/lake floor, they are inferred to continue as topographically controlled sediment gravity flows. These sediments are frequently reported from the geological record and are inferred to be one of the main parts of the growing volcanic edi- fice (WHITE 1996, SMELLIE and HOLE 1997). Identifying the link between the eruption and the resulting deposit is crit- ical to interpreting these sediment gravity flow deposits, especially distinguishing them from other, non-volcanic grav- ity current deposits. Deposits resulting from pyroclastic density currents directly initiated by subaqueous volcanic erup- tions are named eruption-fed pyroclastic density current deposits and are grouped according to their eruption column dynamics (MUELLER and WHITE 1992, WHITE 2000, MUELLER 2003). The behaviour of a subaqueous eruption column inferred to be controlled predominantly by the mass flux of the magma versus the mass fraction of the external water (KOYAGUCHI and WOODS 1996). In the case of a low external water fraction in the eruption column (due to a high mag- matic flux), the column will be steam-poor and therefore dense. Such columns are inferred to be low and probably eas- ily collapse shortly after eruption, forming dense pyroclastic gravity flows that are more or less directly initiated from the crater. If particle concentrations are very high, the resulting deposits may even be welded. In the case of a high water fraction (due to a lower magmatic flux) the column can be charged in steam, and therefore could behave more buoy- antly. The column will grow until it becomes too heavy for rise further, leading to collapse along its margins and spawn- ing column margin-fed density currents (WHITE 2000). If the column hits the water surface, extensive laterally spread- ing mushroom-like clouds may form, from which ash gradually falls out to form graded fallout beds (CASHMAN and FISKE 1991). If the rising column ingests more water it becomes gravitationally unstable and collapses to form low tem- perature, dense pyroclastic gravity flows (WHITE 2000). Formation of eruption columns during shallow subaqueous eruptions is mainly restricted to the quasi-steady state eruption phases, commonly referred as continuous up-rush activ- ity (Figure 8.8) (KOKELAAR 1983, 1986). In vigorous eruptions in subaqueous environments, water exclusion zones can 193 Figure 8.8. A model cartoon of a small-volume shallow-water, Surstseyan-type eruption with continuous uprush con- ditions (shown on the left hand side of the figure “A”) and individual tephra jets (shown on the right hand side of the figure “B”). Volcanic facies of lapilli tuff bed-dominated succession deposited from subaqueous eruption-fed density currents (1) and lapilli tuff breccia-dominated succession rich in ballistically emplaced impact structures formed under a steam cupola (after MUELLER 2003: p. 201, plate 4)194develop near the vent zone that allow formation of fluidal shaped bombs and armoured lapilli; armoured lapilli are gen- erally considered to be able to form only in subaerial conditions, so careful work is required to reconstruct eruption conditions to determine whether eruptions took place entirely underwater or were partly emergent (Figure 8.9) (KOKELAAR 1983, KOKELAAR and DURANT 1983, KOKELAAR and BUSBY 1992, WHITE 1996, WHITE and HOUGHTON 2000). During low magma discharge rates, only intermittent explosions take place and activity is dominated by tephra jetting (WHITE 1996, MASTIN and WITTER 2000, NÉMETH et al. 2006). The deposits of tephra jets are poorly understood. During shallow subaqueous to emergent eruption phases (e.g. Surtla), mostly subaqueous pyroclastic density currents carry fragmented pyroclastic particles away from the vent (WHITE 1996, 2000). However, after explosions breach the water surface, aerial transport begins to dominate; that is, dense eruption columns are formed that generate fallout-dom- inated deposits as well as the deposits of horizontal base surges (LORENZ 1974a, b). The ratio between fall versus base surge dominated deposits is very variable in the resulting tephra cone, as documented from Surtsey (LORENZ 1974a). As a result of continued deposition a tephra cone grows above water level, commonly forming one or a few crescent-shaped or sub-circular islands (SOHN and CHOUGH 1992, WHITE 2001). These islands are very fragile, being made up of loose tephra deposits, and are strongly at the mercy of wave action. In most cases they are short-lived, as in the case of Graham Island, which formed in 1831 just south of Sicily, sparking a three-way international dispute over its ownership before it disappeared beneath the waves eight months later. In rare cases these islands can be efficiently armoured by solid rock if the eruption becomes “dry” and a lava fountain forms in the last phase of the sequence, leading to the formation of a small lava shield in the crater, as at Surtsey (THORARINSSON 1965, 1967, LORENZ 1974a). It has also been reported that the immediate and ongoing palagonitisation of the volcanic glass shards in the fine tephra could form hard and imper- meable beds in the ejecta construct (THORARINSSON 1965, JAKOBSSON 1972). Deep subaqueous explosive volcanism The range of processes responsible for generating explosive eruptions in deep water and transporting and depositing eruption products are poorly understood. The very limited access to sites on active volcanoes and the problem of inter- preting stratigraphical relationships between volcanic and non-volcanic rock units inferred to form in deep subaqueous environments hinder our understanding of this type of volcanism (WHITE et al. 2003). In deep subaqueous environments explosive eruptions driven by magmatic gases can occur (HEAD and WILSON 2003). Hawaiian- and Strombolian-style eruptions are predicted in deep water, and their eruption products are inferred (HEAD and WILSON 2003). The basic idea of such eruptions take place is that magmatic volatiles can accumulate prior an erup- tion, and then the over-pressurized magma can fragment in deep water in a similar way that it would in subaerial condi- tions (HEAD and WILSON 2003). Young deposits interpreted to be result of such deep water hawaiian or strombolian erup- tions have now been identified in many places in the sea floor in water depths ranging from 1000 to 4300 metres (CLAGUE et al. 2003). Because this type of eruption in deep subaqueous environments produces only low eruption columns, only very fine pyroclasts are expected to be transported and deposited extensively on the sea/lake floor, such as “limu o Pele” (limu shells) (MAICHER and WHITE 2001, CLAGUE et al. 2003). Many authors suggest that these deep water eruptions are dominantly driven by magmatic fragmentation, combined with varying degrees of fragmentation driven by magma–water interaction (BATIZA and WHITE 2000). Because steam can theoretically only be produced at water depths shallower than the critical depth of sea water (about 3 km), it is expected that fragmentation by explosive magma–water interaction will be of importance only at depths of 3 km or less. In the case of acidic magmas, eruption of highly vesicular pyroclasts forms pumice, which is capable in floating in water until it becomes water-saturated and sinks as its density exceeds that of water. In addition to explosive eruptions, pumice can form in association with subaerial and subaqueous lava flows (MCPHIE and ALLEN 1992, MANLEY 1996, AKAY and ERDOGAN 2001). However, pumice-rich deposits formed in association with subaqueous lava flows can be car- ried away by currents, and may be deposited far from their source lava flow. In this instance, identification of pumice deposited by sinking of clasts in pumice rafts in subaqueous settings is not unambiguous evidence for nearby explosive subaqueous eruptions (WHITE et al. 2003). Pumice-forming subaqueous eruptions are divided into four categories (Figure 8.10) (KANO 2003); 1) subaqueous Plinian-types, 2) those generating subaqueous flows, 3) those involving explosive 195 Figure 8.9. Schematic model for subaqueous lava fountain eruption (after MUELLER and WHITE 1992). A) Initiation of subaqueous lava fountain on the sea/lake floor and formation of steamy, dense eruption-fed current. B) Eruption plume fully develops when erup- tion progress. Eruption-fed pyroclastic density currents form. Along the eruption cloud margin increased hydroclastic fragmentation occurs. In steam envelop, hot, fluidal pyroclasts can form. C) In the final stage of the eruption the steam envelope breaks, and dilute eruption-fed density currents passing the vent. D) In a theoretical section changes from magmatic to hydromagmatic fragmentation are apparent on the basis of the textural changes of the deposits formed (deposition from high particle concentration density currents to dilute ones). E) Theoretical relationships between the location of clasts formed in the column in the steam envelop and by the pass- ing density currents (from high particle concentration to more dilute ones) bulk interaction of vesiculating magma and water and 4) those ones generating pumice during non-explosive processes. Welding of subaqueous pumice deposits is under debate, and various explanations and conditions are inferred to explain their generation (SPARKS et al. 1980, CAS and WRIGHT 1991, KOKELAAR and BUSBY 1992, KOKELAAR and KONIGER 2000, SIMPSON and MCPHIE 2001). An increasing number of subaqueous calderas have been identified on the seafloor recently on the basis of high res- olution bathymetry (CAS 1992, WRIGHT and GAMBLE 1999, YUASA and KANO 2003). These features have significant eco- nomic value because of their potentially extensive mineralization. Their proximal deposits in particular are potentially 196 Figure 8.10. Summary cartoons of the possible ways to form pumiceous pyroclasts during subaqueous eruptions (after KANO 2003: A) p. 225, fig. 16, B) p. 224, fig. 14, C) p. 223, fig. 13). A) Quench and mechanical fragmentation in the margin of a subaqueous pumiceous dome. Hot pumice blocks ascend to the surface and after drifting settle on the seafloor. Especially larger pumice blocks can fragment further upon contact with sea water. Fragmentation can be enhanced by phreatomagmatic eruptions from the interior of the pumiceous dome as well. B) Subaqueous Vulcanian-style eruption can disrupt the top of a pumiceous dome (I) and generate coarse pumice rich mass flows on the flank of the dome (II). C) A general model to show subaqueous flow-generating processes. “A” represents the gas-supported eruption plume. “B” is the zone where mixing of eruption plume and water takes place. “C” is the zone where buoyancy carries pumice and ash from the convective plume. “D” is the zone where fine pyroclasts transported in sus- pension. “E” is the region where water-logged pumice and dense lithics fall back. “F” represents gas- and water-logged density cur- rents. “G” is the region where gas-supported hot subaqueous pyroclastic flows operate. “H” is the accumulation zone of pumice. D) Hot pumices are charged with gases and buoyant. Until they are hot, steam envelop prevents pumice clasts from hydroclastic frag- mentation, rapid cooling and water-entrance to the vesiclesof economic significance. Since these identified calderas are large (tens of km across) features, it is expected that exten- sive distal pyroclastic successions may be associated with them as well. Subglacial volcanism Subglacial eruptions refer to those eruptions where the erupting vent is located under ice cover (Figure 8.11) (SMELLIE 2000). This could occur in any type of glaciated area, such as high mountains or close to the poles where permanent ice sheets are present (SMELLIE 2000). Volcanoes formed in englacial settings share many features with vol- canoes formed in other subaqueous settings (e.g. in lacustrine environments). This is because the eruption can pro- duce large volumes of melt water that acts as a subaqueous environment below the ice sheet (SMELLIE and SKILLING 1994, TUFFEN et al. 2001, SIGVALDASON 2002, GUDMUNDSSON 2003, SCHOPKA et al. 2006, SMELLIE 2006, STEVENSON et al. 2006). Eruptions under ice can generate large volumes of hyaloclastite and melt water, which can initiate floods termed jökulhlaup (glacier outburst floods). Jökulhlaup are devastating floods capable of depositing large volumes of sediment and causing significant landscape modification (CAREY et al. 2000, ALHO 2003, BJORNSSON 2003, CARRIVICK et al. 2004, ROBERTS 2005, EVA T T et al. 2006, RUSHMER 2006, RUSSELL et al. 2006, SMITH et al. 2006). Their power is well demonstrated in many Icelandic subglacial eruptions, such as the October 1996 eruption beneath the Vatnajökull glacier (GUDMUNDSSON et al. 1997, MARIA et al. 2000, GUDMUNDSSON et al. 2004, STEFANSDOTTIR and GISLASON 2005). Eruptions beneath glaciers are largely controlled by the thickness of the ice cover, e.g. thin (100–150 m) or thick (over 150 m) (SMELLIE and SKILLING 1994, SMELLIE 2000). The classification of the resulting volcanic structure is commonly based on the relative proportion of volcanic, volcaniclastic, and sedimentary lithologies present, perhaps combined with their 3D architecture (Figure 8.11) (SMELLIE 2000). 197 Figure 8.11. General model of subglacial eruptions and the formation of table mountains with the associ- ated volcanic facies distribution (KOMATSU et al. 2007: fig. 5)198 Figure 8.13. Theoretical model illustrating the environmentally controlled eruptive processes of an active englacial volcano in a narrow ice-bound lake (after WERNER and SCHMINCKE 1999: p. 350, fig. 10A–D). A) Effusive/intrusive deep water eruption and accumulation of pillow lavas. B) Hydroclastic fragmentation and shallow water, Surtseyan-style eruption. C) Emergent stage, more magmatic fragmenta- tion (Strombolian- and/or Hawaiian-style eruption). D) Effusive emergent stage producing lava flows over the volcanic edifice Figure 8.12. A model architecture of a tuya (CHAPMAN 2002: p. 278, fig. 5) In the case of thin ice cover, the eruption quickly generates a cavity (vault) in the ice and melt water escapes through a complex network of tunnels generated by the heat of the magma (SMELLIE 2000). The eruption site quickly becomes open to air and phreatomagmatic eruptions triggered by magma–water interaction often forms a tuff cone in the newly formed hole in the glacier. The vent site of this new phreato- magmatic volcano is quickly sealed off from external water, and eruptions turn toward ‘drier’lava fountaining hawaiian- or mildly explosive Strombolian-style events. In the final stages the heat of the newly formed volcano may melt more ice at the margins of the ice hole, and melt water can be dammed between the newly formed cone and the ice vault walls. Sub- sequent eruptions pour lava flows into this meltwater lake, forming lava deltas and associated hyaloclastite units. During eruptions through thicker ice piles, the initially formed melt water can be sealed and locked into the newly formed cavity over the vent, creating prefect subaqueous con- ditions for the subsequent eruption(s) (SMELLIE 2000). This condition can be reached especially in those glaciers where the upper permeable layer of the glacier is thick. In this case the melt water depth is not sufficient to lift up the ice cover by floating, and therefore melt water cannot leave the erup- tion site other than by slow escape along thermally eroded channels. When lava effusion takes place, lava can enter the englacial lakes and accumulate thick piles of hyaloclastite, preserved as table mountains (Figure 8.12) (WERNER et al. 1996, WERNER and SCHMINCKE 1999) (stapi in Icelandic, or tuyas in British Columbia) after the ice cover disappears. Before the growing volcanic pile can emerge and become fully subaerial a palagonite tuff and tuff breccia composed edifice develops, commonly referred to as tindar (SMELLIE 2000). Explosive eruptions that take place in the trapped melt water are typical of those in shallow subaquous environments (Figure 8.13). 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Pillow lava cross section from a sublacus- trine lava flow from the Western Snake River Plain, Idaho, USA. Note the typical radial joint pattern of the pillow lava and the brownish rim composed of palagonitized glassy lava. 4. Overview of the Pálháza hyaloclastite succession surrounding irregular shape coherent rhyolite intrusive/extrusive bodies. 5. A tuff cone formed in the Lake Voui caldera lake on Ambae Island, Vanuatu in December 2005. Dark tephra jets erupt from the vent.‘A’a: Hawaiian word used to describe a lava flow whose surface is broken into rough angular fragments. Accessory: A mineral whose presence in a rock is not essential to the proper classification of the rock. Accessory fragment: A lithic fragment composed of country rock that has been explosively ejected during an eruption (CAS and WRIGHT 1987: p. 54). Accessory fragments within pyroclastic deposits may be difficult to distinguish from accidental fragments. In general terms, referred to as a xenolith. Accidental: Pyroclastic rocks that are formed from fragments of non-volcanic rocks or from volcanic rocks not related to the erupting vol- cano. Accidental fragment: A clast picked up locally by pyroclastic flows and surges (CAS and WRIGHT 1987: p. 54). Accidental fragments may be difficult to distinguish from accessory fragments. In general terms, referred to as a xenolith. Accidental rock fragments in maar and tuffring deposits provide additional important information: (1) their maximum size allows estimation of explosion energy; (2) the type of crustal rocks present permits inferences about explosion depths if crustal stratigraphy is known; and (3) mantle-derived ultramafic xeno- liths, are common in some maar deposits. Acid: A descriptive term applied to igneous rocks with more than 60% silica (SiO 2 ). Accretionary lapilli: Accretionary lapilli are mud balls which result from a wet nucleus falling through a volcanic ash cloud. They flatten on striking the ground or may roll on loose ash and grow like a snowball. Spherical aggregates (commonly with a concentric structure) formed by the accretion of moist ash in eruption clouds (WHITE and HOUGHTON 2000: p. 495). Also used for all ash aggregates, includ- ing mud lumps (HOUGHTON et al. 2000: p. 513). Achnelith: A type of juvenile fragment characterized by smooth, glassy molded surfaces formed from lava spray from extremely fluid mafic eruptions (WALKER and CROASDALE 1972). Agglomerate: Agglomerate is a mix of volcanic material which has been solodified into a rock. A course, pyroclastic deposit composed of a large proportion of fluidal-shaped volcanic bombs that are formed, in the strictest sense, by a fall deposit in the immediate vicinity of a volcanic vent. It is best applied to describe bomb and scoria deposits that build strombolian cones, and should never be used as a non- generic term for a “volcanic breccia” (CAS and WRIGHT 1987: p. 359). Agglutinate: A pyroclastic deposit consisting of an accumulation of originally plastic ejecta and formed by the coherence of the fragments upon solidification. Agglutination: Instantaneous flattening of hot, soft pyroclasts upon landing. The resultant deposit is an agglutinate or spatter pile; particle outline is in part retained (WOLFF and SUMNER 2000: p. 321). Airfall: Volcanic ash that has fallen through the air from an eruption cloud. Airfall deposits are characteristically well-sorted and well-lay- ered, and typically exhibit mantle bedding (CAS and WRIGHT 1987). Alert level: Alert level is a measure of the current status of the volcano. Amygdaloidal: A volcanic texture comprising vesicles (rounded holes resulting when magma cools around gas bubbles) which have been sub- sequently filled by secondary minerals. Amygdule: An individual vesicle which has been subsequently filled-in by secondary minerals. Annular flow: One of four two-phase flow regimes, in which magma lines the conduit walls and gas flows in a central jet (JAUPART 2000: p. 237). Armoured lapilli: A type of accretionary lapilli composed of a crystal, pumice, or lithic fragment core which is surrounded by a rim of fine to coarse ash (MCPHIE et al. 1993: p. 29). Ash: A textural term for volcanic fragments less than 2 mm in diameter (FISHER 1966; SCHMID 1981). Ash is the typical product of explosive volcanic eruptions. Measuring less than 1/10 inch in diameter, ash may be either solid or molten when first erupted. By far the most com- mon variety is vitric ash (glassy particles formed by gas bubbles bursting through liquid magma). Ash cloud: A cloud of ash produced during pyroclastic eruptions. These clouds can result from rapid rising of the hot, buoyant ash-rich erup- tive plume, or can be derived by elutriation at the top of a pyroclastic flow (CAS and WRIGHT 1987). Ash flow: A type of pyroclastic flow comprising dominantly ash-sized particles. Hot ash flows may be called “glowing avalanches” or “nuee ardentes”, and if their volume is large enough, may eventually form deposits known as welded tuffs. These types of flows are extremely dangerous and historically have killed hundreds of thousands of people. A turbulent mixture of gas and rock fragments, most of which are ash-sized particles, ejected violently from a crater or fissure. The mass of pyroclastics is normally of very high temperature and moves rapidly down the slopes or even along a level surface. 207Ashfall (Airfall): Volcanic ash that has fallen through the air from an eruption cloud. A deposit so formed is usually well sorted and layered. Atmospheric shock wave: A strong compressional shock wave caused by a combination of volcanic ejecta and sonic waves. Autoclastic volcanic breccias: Result from internal processes acting during movement of semisolid or solid lava; they include flow breccia and intrusion breccia (FISHER 1960). Avalanche: A large mass of material or mixtures of material falling or sliding rapidly under the force of gravity. Avalanches often are classi- fied by their content, such as snow, ice, soil, or rock avalanches. A mixture of these materials is a debris avalanche. Ballistic projectiles: Ballistic projectiles are pieces of rock thrown from a volcanic vent in an eruption. They are generally confined to less than 3 kilometres radius from the vent because of their size. Base surge: Base surges are ground hugging, outward moving clouds of gas and ash. They result from water–magma interactions. A turbu- lent, low-density cloud of rock debris, water, and/or steam that moves over the ground surface at extremely high speeds. Base surges are commonly the result of directed volcanic explosions. Base surge deposits are commonly composed of cross-bedded deposits comprising ash and lapilli. During 1965 Taal volcano (Philippines) erupted and base surges travelled 4 km and killed 189 people. Base surges were first identified during ocean nuclear weapons explosions in the Pacific. Bedding sags: Also known as “bomb sags” (WENTWORTH 1926), form by the impact of ballistically-ejected bombs, blocks and lapilli into beds capable of being plastically deformed. They are characteristic of hydroclastic deposits and have been described from the deposits of many maar volcanoes, tuff rings and tuff cones. Beds beneath the fragments may be completely penetrated, dragged down and thinned, fold or show micro-faulting (HEIKE ~1971). Deformation is commonly asymmetrical, showing the angle and direction of impact if three-dimen- sional exposures are available. Bed forms: The surface configuration of a bed (V ALENTINE and FISHER 2000: p. 571). Bed forms by base surges: Bed forms occur as three main kinds — sandwave, massive and planar (plane parallel) beds (SCHMINCKE et al. 1973; SHERIDAN and UPDIKE 1975), and are grouped into three facies types (WOHLETZ and SHERIDAN 1979) related to a fluidization model of transport and deposition. FISHER and WATERS (1970), FISHER and CROWE (1973) and SCHMINCKE et al. (1973) have emphasized bed forms in terms of the flow regime concept These different approaches are treated separately although they are not mutually exclusive. Bedset: A sequence of beds with distinct internal structures, textures, colours or compositions that sets them apart from other sequences, usu- ally bounded by unconformities, or by fallout layers (VALENTINE and FISHER 2000: p. 571). Bench: The unstable, newly-formed front of a lava delta. Blister: A swelling of the crust of a lava flow formed by the puffing-up of gas or vapor beneath the flow. Blisters are about 1 meter in diam- eter and hollow. Block: Angular chunk of solid rock ejected during an eruption. Fragments of solid rock greater than 64 millimeters in diameter that are eject- ed during volcanic eruptions. Blocks are commonly composed of accessory fragments made up of crystallized magma associated with the eruption (e.g. pieces of a lava dome). Block-and-ash flow deposit: Small-volume pyroclastic flow deposit characterized by a large fraction of dense to moderately vesicular juve- nile blocks in a medium to coarse ash matrix of the same composition (FREUNDT et al. 2000: p. 581). Blocky lava: Lava flows that are characterized by highly fractured surfaces which contain fragments of debris (usually flow fragments) up to several metres in diameter. The size of the surface fragments in blocky lavas is controlled by the rheology of the lava in the interior of the flow (KILBURN 2000: p. 291). Bomb: Fragment of molten or semi-molten rock, 2 1/2 inches to many feet in diameter, which is blown out during an eruption. Because of their plastic condition, bombs are often modified in shape during their flight or upon impact. Bubbly flow: One of four two-phase flow regimes, in which the gas phase appears as bubbles suspended in a continuous magma phase (JAUPART 2000: p. 237). Bulking: The erosion and incorporation of secondary, exotic debris by lahars as they move downstream (V ALLANCE 2000: p. 601) Caldera: The Spanish word for cauldron, a basin-shaped volcanic depression; by definition, at least a mile in diameter. Such large depres- sions are typically formed by the subsidence of volcanoes. Crater Lake occupies the best-known caldera in the Cascades. Caldera cycle: A commonly observed evolutionary sequence recognized in many caldera complexes. From oldest to youngest, the seven stages of the caldera cycle are: 1) regional tumescence and generation of ring fractures; 2) ignimbrite (pyroclastic) eruption(s); 3) caldera collapse; 4) pre-resurgent volcanism and intra-caldera sedimentation; 5) resurgent doming; 6) major ring fracture volcanism; and 7) ter- minal fumarolic and/or hot spring activity. Capping stage: Refers to a stage in the evolution of a typical Hawaiian volcano during which alkalic, basalt, and related rocks build a steeply, sloping cap on the main shield of the volcano. Eruptions are less frequent, but more explosive. The summit caldera may be buried. Cinder cone: A volcanic cone built entirely of loose fragmented material (pyroclastics). Clastogenic lava: A lava flow formed by the rheomorphic flow of coalesced and agglutinated hot pyroclasts, typically fed by a lava fountain (WOLFF and SUMNER 2000: p. 321). Coalescence: The process by which hot fluidal pyroclasts form a homogeneous liquid in which the particle outlines are obliterated (WOLFF and SUMNER 2000: p. 321). Cognate lithic fragment: Non-vesiculated juvenile magmatic fragments that have silicified from the erupting magma (CAS and WRIGHT 1987: p. 54). Column collapse: Column collapse is caused by an eruption column reaching a critical level then collapsing under its own weight. This forms the most dangerous of volcanic processes — pyroclastic flows and surges. Columnar jointing: A type of fracture pattern resulting from the thermal contraction of hot volcanic rocks after their crystallization which commonly is expressed in elongate, pentagonal or hexagonal columns oriented perpendicular to the cooling surface. Columnar jointing is common in all compositions of lava flows, although it is generally best developed in mafic (basalt) lava flows and in felsic welded tuffs. Continuous uprush: A style of explosive eruption in shoaling volcanoes, which combine into a continuous uprush and are characterized by a tall non-spreading eruption column (SMELLIE 2000: p. 403). Cooling unit: A group of hot pyroclastic deposits (ignimbrites) that cools at more or less the same time. A deposit from a single eruption that shows simple variations in the degree of welding is known as a simple cooling unit. When many ignimbrites occur over an extremely short period of time, each individual ignimbrite may be deposited, and start to weld over a previous deposit or group of deposits that are cool- ing and undergoing welding. The resulting deposits have several zones of partial and dense welding, and since they more or less cool together, are known as compound cooling units (CAS and WRIGHT 1987: p. 253–255). 208Coulée: A type of rhyolite lava flow that forms when lava issues from one side of a volcanic vent and produces a lava flow which is elongate in plan view (CAS and WRIGHT 1987: p. 81). Composite volcano: Relatively large, long-lived constructional volcanic edifice, comprising lava and volcaniclastic products erupted from one or more vents, and their recycled equivalents. Compound volcano: Volcanic massif formed from coalesced products of multiple, closely spaced, vents. Conduit: A passage followed by magma in a volcano. Country rocks: The rock intruded by and surrounding an igneous intrusion. Crater: A steep-sided, usually circular depression formed by either explosion or collapse at a volcanic vent. Cryptodome: A cryptodome is a mound caused by the accumulation of viscous magma just beneath the surface. Cupola (Water dome): A water cupola is a dome formation on the surface of the water just before an underwater eruption breaks through. The dome may rise to form a cylinder. Cupolas have been observed with a base of 100-200 m and height of 26 m. Curie point: The temperature at which a body loses (by heating) or preserves (by cooling) its permanent magnetization. As rocks cool, the electromagnetic field aligns magnetic minerals in the magma, and their orientation is preserved as the rocks cool below the Curie point. Curtain of fire: A row of coalescing lava fountains along a fissure; a typical feature of a Hawaiian-type eruption. Debris avalanche: A rapid and unusually sudden sliding or flowage of unsorted masses of rock and other material. Catastrophic landsliding of gravitationally unstable volcano flanks resulting in a widely dispersed deposit at the foot of the edifice, typically characterized by a hummocky surface. As applied to the major avalanche involved in the eruption of Mount St. Helens, a rapid mass movement that includ- ed fragmented cold and hot volcanic rock, water, snow, glacier ice, trees, and some hot pyroclastic material. Most of the May 18, 1980 deposits in the upper valley of the North Fork Toutle River and in the vicinity of Spirit Lake are from the debris avalanche. Debris ava- lanches differ from debris flows in that they are not water saturated and in that the load is entirely supported by particle-particle interac- tion (VALLANCE 2000: p. 601). Debris flow: A mixture of water-saturated rock debris that flows downslope under the force of gravity (also called lahar or mudflow). A type of mass flow comprising a dense, cohesive, flowing mixture of sediment (mud through boulder sized materials, generally >50% by vol- ume), water, and commonly, organic debris. Debris flows generally move downslope in laminar fashion due to the force of gravity (VALLANCE 2000: p. 601; CAREY 2000: p. 627). Debris flows generated at volcanoes are commonly referred to as lahars. A uniform mix- ture of solid and liquid phases in vertical profiles characterizes debris flows and distinguishing them from more water-rich hyperconcen- trated flows (VALLANCE 2000: p. 601). Debris fall: A debris fall is the near free fall of debris from an overhang or vertical face. Debulking: A process in which the lahar selectively deposits certain particles, owing to their size or density, as it moves downstream (VALLANCE 2000: p. 601). Decompressive melting: Melting that occurs when rocks undergo a decrease in pressure. This commonly occurs in the vicinity of hot spots as mantle rocks rise to shallower levels in the earth due to convective rise and upwelling (SIGURDSSON 2000: p. 15). Melting occurs as a result of decreasing pressure, not increasing temperature Destructiveness index: the logarithm of the area covered by lava, pyroclastic flows, and surges, or buried under more than 100 kg/m 2 of tephra during an eruption (PYLE 2000: p. 263). Devitrification: The solid-state transformation of volcanic glass into crystalline materials. Devitrification tends to be more prevalent in dense- ly-welded tuffs, but may also occur in less densely-welded or unwelded pyroclastic and/or volcaniclastic deposits. The main products of devitrification are cristobalite (SiO 2 ) and alkali feldspar (KAlSi 3 O 8 ) (CAS and WRIGHT 1987: p. 258). Diatreme: A funnel-shaped, pipe-like volcanic conduit, usually filled with volcaniclastic debris, emplaced by the explosive energy of gas- charged magmas. Diatremes are believed to result from hydrovolcanic fragmentation and subsequent wall rock collapse (VESPERMANN and SCHMINCKE 2000: p. 683), and may reach depths up to 2500 metres. Diamond-bearing diatremes are economically important and are referred to as kimberlite pipes. Dike: A discordant, sheetlike body igneous body formed from the injection of magma into a fracture within the brittle crust of the earth (CARRIGAN 2000: p. 219; MARSH 2000: p. 191). Generally a tabular igneous body which cross-cuts the planar structures in the adjacent rocks. Directed blast: A hot, low density mixture of gas, rock debris, and ash that is propelled by a volcanic eruption and generally moves along the ground at high speeds. Dispersed flow: One of four two-phase flow regimes, in which magma takes the form of fragments in a continuous gas phase (JAUPART 2000: p. 237). Dispersal index (D): A measure of the area covered by a pyroclastic fall deposit, specifically the area enclosed by an isopach draw at 1/100 of the maximum thickness of the deposit (HOUGHTON et al. 2000: p. 513). Dome: A steep-sided mass of lava that is generally formed immediately above the volcanic vent from which it was extruded. Domes are gen- erally circular in plan and have a relatively small surface area relative to other types of lava flows. Domes may be spiny, rounded, or flat on top, and often have rough, blocky surfaces formed by the fragmentation of the dome’s crust during intrusion. Domes may grow by extrusion of lava onto the outer surface of a previously formed dome (exogenous dome) or may be formed by inflation of a pre-existing dome (endogenous dome). Domes are most commonly the result of extrusion of viscous lava (primarily of the composition of rhyolite and dacite, but andesite may occur as well). Downsag caldera: A type of caldera characterized by inward sloping topography, inward tilted wall rocks, and an apparent absence of large displacement caldera bounding faults (LIPMAN 1997). Downsag calderas are believed to result from small volume eruption from a deep- seated subvolcanic intrusion. Edifice (Volcanic): A volcanic edifice is a constructional feature built from erupted material. Ejecta: Material that is thrown out by a volcano, including pyroclastic material (tephra) and lava bombs. Elutriation: Loss of small particles by an upward flow of gas through a deposit. Epiclastic volcanic breccias: Result from transportation of loose volcanic material by epigenetic geomorphic agents, or by gravity, and include laharic breccia, water-laid volcanic breccia, and volcanic talus breccia (FISHER 1960). Episode: An episode is a volcanic event that is distinguished by its duration or style. Eruption: The process by which solid, liquid, and gaseous materials are ejected into the earth’s atmosphere and onto the earth’s surface by vol- canic activity. Eruptions range from the quiet overflow of liquid rock to the tremendously violent expulsion of pyroclastics. 209210 Eruption cloud: The column of gases, ash, and larger rock fragments rising from a crater or other vent. If it is of sufficient volume and velocity, this gaseous column may reach many miles into the stratosphere, where high winds will carry it long distances. A volcanic cloud is a convoluted rolling mass of water vapour and ash that is highly charged with electricity and overhangs a volcano during an eruption. The cloud is produced by a column of gases, ash, and rock emitted from a crater. Eruption clouds may reach great heights. The 1883 eruption of Krakatau produced an eruption column 50 km high. Eruption clouds may take on the shape of pine trees or cau- liflowers. Eruption column: An eruption column is the lower vertical part of the eruption cloud, where the ash and gases rise with great speed. Eruptive vent: The opening through which volcanic material is emitted. Extrusion: The emission of magmatic material at the earth’s surface. Also, the structure or form produced by the process (e.g. a lava flow, volcanic dome, or certain pyroclastic rocks). Facies: A part of a rock body that can be differentiated from another part of a related rock body by textural or compositional variations. The general appearance or composition of one part of a rock body as contrasted with other parts (AGI 1976: p. 155). Facies changes: The textural and compositional changes that occur laterally and/or vertically within related rock bodies. Fall deposit: Mantling blanket of pyroclastic particles (ash, scoria, pumice, etc) erupted explosively and transported through the atmosphere before falling back to the ground. Felsic: An igneous rock having abundant light-colored minerals. Fire fountain: See also: lava fountain Fissures: Elongated fractures or cracks on the slopes of a volcano. Fissure eruptions typically produce liquid flows, but pyroclastics may also be ejected. Flank eruption: An eruption from the side of a volcano (in contrast to a summit eruption). Flood basalt: voluminous and thick sequences of mafic lavas erupted from fissures over relatively short periods of time (plateau basalts) (PEFIT and DA VIDSON 2000: p. 89). Flow banding: A foliation commonly observed in intermediate and felsic lavas, that results from shearing of the lava during laminar flow (CAS and WRIGHT 1987: p. 78). In rhyolite flows, flow banding is commonly exhibited by alternating bands comprising volcanic glass and spherulites (small, radiating bodies of devitrified glass). Flow regime: Hydraulic conditions of noncohesive flow of sand and silt that develop ripples, dunes, plane parallel beds, and antidunes. Progressive changes in bed forms occur as flow regime increases. Low-flow regime conditions form small scale ripples that progress to dunes; high flow regime conditions form plane beds and then antidunes (VALENTINE and FISHER 2000: p. 571). Flow transformation: Reversable changes in sediment gravity flows between turbulent and steady flow related chiefly to particle concentra- tion, thickness of flow, and flow velocity (VALENTINE and FISHER 2000: p. 571). Formation: A body of rock identified by lithic characteristics and stratigraphic position and is mappable at the earth’s surface or traceable in the subsurface. Fragmentation: Fragmentation is a process whereby bursting bubbles tear apart magma as it reaches the surface. At a depth of 10 km gas filled magma is at a pressure of 3000 atmospheres. Therefore a large pressure drop occurs as the magma rises towards the surface through the conduits. Fragmentation also can be defined as a transition from a continuous melt with a dispersed gas phase to disconnected parcels of bubbly melt within a continuous gas phase (CASHMAN et al. 2000: p. 421). Fragmentation index: a parameter measuring the grain size of a pyroclastic fall deposit, specifically the percentage of ash finer than 1 mm at the point on the dispersal axis to1/10 of the maximum thickness of the deposit (HOUGHTON et al. 2000: p. 513). Fuel-coolant interaction: The interaction of magma (fuel) with external water (coolant) that may result in thermal explosions (VESPERMANN and SCHMINKE 2000: p. 683). Fumarole: A vent which releases volcanic gases. These include steam (H 2 O), carbon dioxide (CO 2 ), sulphur dioxide (SO 2 ), hydrogen sul- phide (H 2 S), as well as other volatile gases emitted from subterranean magmas. Geyser: Most geysers are hot springs that episodically erupt fountains of scalding water and steam. Such eruptions occur as a consequence of groundwater being heated to its boiling temperature in a confined space (for example, a fracture or conduit). A slight decrease in pressure or an increase in temperature will cause some of the water to boil. The resulting steam forces overlying water up through the conduit and onto the ground. This loss of water further reduces pressure within the conduit system, and most of the remaining water suddenly con- verts to steam and erupts at the surface. Guyot: A type of seamount that has a platform top. Named for a nineteenth-century Swiss-American geologist. Hawaiian eruption: Hawaiian eruptions describe a style of eruption typically seen on shield volcanoes. Hawaiian eruptions involves lava fountains and lava flows. Commonly seen in Hawaii, Iceland, Reunion, Hawaiian eruptions are spectacular and the most photogenic of all eruption types. Heat transfer: Movement of heat from one place to another. Horizontal blast: An explosive eruption in which the resultant cloud of hot ash and other material moves laterally rather than upward. Hornito: A small rootless spatter cone that forms on the surface of a basaltic lava flow (usually pahoehoe) is called a hornito. A hornito devel- ops when lava is forced up through an opening in the cooled surface of a flow and then accumulates around the opening. Typically, hor- nitos are steep sided and form conspicuous pinnacles or stacks. They are “rootless” because they are fed by lava from the underlying flow instead of from a deeper magma conduit. Hot spot: A volcanic center, 60 to 120 miles (100 to 200 km) across and persistent for at least a few tens of million of years, that is thought to be the surface expression of a persistent rising plume of hot mantle material. Hot spots are not linked to arcs and may not be associat- ed with ocean ridges. Hot-spot volcanoes: Volcanoes related to a persistent heat source in the mantle. Hummocks: These are rounded or conical mounds within a volcanic landslide or debris avalanche deposit. Hummocks contain a wide range of rock debris, reflecting the variation of deposits that previously formed the flanks of the volcano. Some hummocks contain huge intact blocks tens to hundreds of metres in diameter. Some of the original layering of lava flows and other deposits can be seen in these large hummocks, but most of the large rock fragments are thoroughly shattered. In other hummocks the rock debris is thoroughly mixed as if the material had been in a blender and thoroughly mixed together. Hyaloclastite: A deposit comprising small, angular glass fragments formed by nonexplosive shattering of lava or magma flowing into water, ice, or water-saturated sediment (BATIZA and WHITE 2000: p. 361; SCHMIDT and SCHMINCKE 2000: p. 383). Hydroclastic rocks: a general term applied to volcanic rocks formed by fragmentation of magma in the presence of water (BATIZA and WHITE 2000: p. 361). Hydrovolcanic eruptions: A general term for eruptions caused by the mixing of magma with water (VESPERMANN and SCHMINKE 2000: p. 683). Encompasses hydroclastic, hydromagmatic, and phreatomagmatic eruptions. Hypabyssal: A shallow intrusion of magma or the resulting solidified rock. Hyperconcentrated flow: A gravitationally driven, nonuniform mixture of debris and water content larger than that of debris flow but less than that of muddy streamflow (VALLANCE 2000: p. 601). Ignimbrite: Fiery raincloud rocks. The rock formed by the widespread deposition and consolidation of ash flows and Nuees Ardentes. The term was originally applied only to densely welded deposits but now includes non-welded deposits. Ignimbrite deposits are poorly sort- ed pyroclastic deposits consisting of glass shards, crystals and lithic fragments. Ignimbrites are formed by the deposition of hot, rapidly expanding, turbulent magmatic gases. Ignimbrites have a volume of 1 cubic km to 2000 cubic km. A term used for pyroclastic flow deposits, that is synonymous with “ash tuff” (LIPMAN 2000: p. 643). According to CAS and WRIGHT (1987: p. 98), the term should only be used to describe pumiceous pyroclastic flow deposits. Isopach: Line joining points of equal thickness of deposit (HOUGHTON et al. 2000: p. 555). Isopleth: Line joining points where the sizes of the largest clasts are the same (HOUGHTON et al. 2000, p. 555). Jökulhlaup: A Jökulhlaup is a glacial outburst caused by melt water from a subglacial volcano. Examples of Jökulhlaup Grimsvötn (Iceland) 1996, 1937. Jigsaw cracks: These are characteristic joint patterns within a debris-avalanche block (UI et al. 2000: p. 617). Juvenile fragment: Glassy or partially crystallized fragments which represent samples of an erupting magma. These include fragments such as pumice, scoria, reticulate, achneliths (Pele’s tears, Pele’s hair), and various types of volcanic bombs (CAS and WRIGHT 1987: pp. 47–53). Kimberlite: Kimberlite is a rock formed by explosive eruptions in volcanic pipes. It is formed at depths of 100-200 km and pressures of 40,000 to 50,000 atm. Kimberlite is the main source of diamonds. Kipuka: Is a Hawaiian term for an “island” of land completely surrounded by one or more younger lava flows. A kipuka forms when lava encircles a hill or a slight rise in the ground as it moves downslope or across relatively flat ground. Because they are surrounded by more recent flows, kipukas are often covered with mature vegetation. Laccolith: A body of igneous rocks with a flat bottom and domed top. It is parallel to the layers above and below it. Lahar: The Indonesian term for a debris flow or a mudflow originating on a volcano (HARRIS 2000: p. 1301). Lahars are generally composed of volcanic materials, but can contain significant amounts of non-volcanic materials derived from erosion during flow. Most volcanolo- gists prefer this term to be used for the process and not the sedimentary deposits that it forms, but unfortunately, this distinction has been largely ignored in the geological literature. Many lahars are composed of sand and coarser materials, and thus, can be distinguished from “mudflows” which predominantly contain silt- or clay-sized grains (RODOLFO 2000: p. 973). A torrential flow of water-saturated volcanic debris down the slope of a volcano in response to gravity. Lahars are also referred to as volcanic mudflows or debris flows. They form in a variety of ways, chiefly by the rapid melting of snow and ice by pyroclastic flows, intense rainfall on loose volcanic rock deposits, break- out of a lake dammed by volcanic deposits, and as a consequence of debris avalanches. Laminar flow: flow regime where viscous effects dominate and flow trajectories are parallel (JAUPART 2000: p. 237). Lapilli: Literally, “little stones.” A textural term for fragments in volcanic rocks and volcanic deposits that range from 2 mm to 64 mm in diameter (FISHER 1966; SCHMID 1981). Lateral blast: A volcanic eruption which is directed horizontally instead of vertically. Lateral blasts may be caused by sudden decompres- sion of a shallow magma chamber residing within the flanks of a volcano (for example, the 1980 eruption of Mt St. Helens), or along the base or side of a lava dome (for example, the 1902 eruption of Mt Pelee in Martinique) (NAKADA 2000: p. 945). Lava: Magma which has reached the surface through a volcanic eruption. The term is most commonly applied to streams of liquid rock that flow from a crater or fissure. It also refers to cooled and solidified rock. An outpouring of lava from a vent or fissure that spreads along the ground surface, as well as the crystallized rock resulting from solidification of the outpouring (PETERSON and TILLING 2000: p. 957). Lava breaching: Breaching is a term to describe a lava flow breakout. The conditions from breaching require a high channel pressure and more fluid lava in the channel compared to the edges. Aa lava is the most prone to breaching. Breaching usually occurs in the upper half of the lava flow. Therefore breaching tends to widen the lava flows rather than extend them length ways. Note: It is rare for breached lava flows to extend more than 50% more than the original flow length. Note: lava flows cool quicker on steep slopes and therefore stops quick- er. Lava delta: Lava entering the sea often builds a wide fan-shaped area of new land called a lava delta. Such new land is usually built on slop- ing layers of loose lava fragments and flows. On steep submarine slopes, these layers of debris are unstable and often lead to the sudden collapse of lava deltas into the sea. Lava dome: Mass of lava, created by many individual flows, that has built a dome-shaped pile of lava. Lava flow: An outpouring of lava onto the land surface from a vent or fissure. Also, a solidified tongue like or sheet-like body formed by out- pouring lava. Lava fountain: A rhythmic vertical fountain-like eruption of lava. Lava lake (Pond): A lake of molten lava, usually basaltic, contained in a vent, crater, or broad depression of a shield volcano. Lava takes are large volumes of molten lava, usually basaltic, contained in a vent, crater, or broad depression. Scientists use the term to describe both lava lakes that are molten and those that are partly or completely solidified. Lava lakes can form (1) from one or more vents in a crater that erupts enough lava to partially fill the crater; (2) when lava pours into a crater or broad depression and partially fills the crater; and (3) a top a new vent that erupts lava continuously for a period of several weeks or more and slowly builds a crater higher and higher above the surrounding ground. A region typically within the summit of a shield volcano which contains partially crystallized or molten lava which lies immediately above a volcanic conduit which joins the lava lake to the magma chamber. Strong magma convection within vol- canic conduits sustains lava lakes within their respective volcanic vents (WALKER 2000: p. 285). Lava levee: A lava levee is a boarder of lava blocks at the edge of a flow. Lava shields: A shield volcano made of basaltic lava. Lava tube: A tunnel formed when the surface of a lava flow cools and solidifies while the still-molten interior flows through and drains away. Limu: Limu, or Limu o Pele (Hawaiian for “seaweed of Pele”), consists of thin flakes of basaltic glass that sometimes form when pahoehoe lava pours into the ocean. As waves wash atop exposed streams of lava, some water may become trapped and boil, resulting in delicate 211steam-filled bubbles of lava. Abrupt chilling and continued expansion of the delicate bubble walls form thin plates and shattered pieces of brownish-green to nearly-clear glass. Littoral cone: A cone of lava fragments built on the surface of a lava flow pouring into a body of water, usually the sea, is called a littoral cone (“littoral” refers to a shoreline). Lava entering the ocean heats and boils seawater, often generating steam explosions that hurl tephra onto the shore, including spatter, bombs, blocks, ash, lapilli, and, rarely, limu. As the various tephra accumulates on the shoreline, a well- developed cone may be created. Lithophysae: Radial aggregates of fibrous crystals which have formed around an expanding vesicle in a melt which is capable of flowing (CAS and WRIGHT 1987: p. 84). Lithophysae are commonly the result of vapor-phase crystallization within a rhyolitic magma. They should not be confused with spherulites, which are similar-shaped structures formed from devitrification of volcanic glass. Lithic: Fragments of previously-formed rocks or dense fragments that occur within volcaniclastic deposits. Lithic fragments may be accesso- ry fragments, accidental fragments, or juvenile fragments. Littoral: An adjective describing physical features or processes associated with shorelines of oceans, seas, or lakes (PETERSON and TILLING 2000: p. 957). Lobate lava: A submarine lava comprising elongate, flattish lobes with smooth, outer glassy skins (BATIZA and WHITE 2000: p. 361). Maar: A maar is a low-relief, broad volcanic crater formed by shallow explosive eruptions. The explosions are usually caused by the heating and boiling of groundwater when magma invades the groundwater table. Maars often fill with water to form a lake. A maar is a type of tuff ring which has been affected by sagging so that it lies below the the level of the surrounding surface. A maar is often filled with water and surrounded by a rim of ejected material that was probably formed by explosive interaction of magma and groundwater. A type of monogenetic volcano, generally formed by subterranean phreatic or phreatomagmatic eruptions that occur as magma explosively inter- acts with ground water or subsurface moisture. Maar craters are cut into the surrounding country rock, vary from 10-500 metres deep, and range from a few hundred metres to 3 km in diameter. Maar volcanoes are generally surrounded by low, shallowly outward-dipping beds of well-bedded volcanic ejecta that rapidly decrease in thickness away from the vent. The volcanic deposits are mainly emplaced by base surges and fallout, and commonly contain very little (or in the case of phreatic eruptions, no) juvenile volcanic materials (VESPERMANN and SCHMINCKE 2000: p. 685; CAS and WRIGHT 1987: pp. 376–377). Examples of Maars Lake Nyos (Cameroon), Suoh (Sumatra, Indonesia), Karpinsky Group (Kurile Islands), Ukinrek Maars (Alaska). Mafic: A compositional term for igneous rocks which contain 45–55% SiO 2 (by weight). Mafic rocks are generally dark coloured, and are characterized by mineralogy including pyroxene and calcium-rich plagioclase, variable amounts of olivine, and accessory minerals such as ilmenite and magnetite. Examples of mafic rocks include basalt and gabbro. Mantle bedding: Pyroclastic deposits generated by ash fall which maintain a uniform thickness and drape over all but the steepest topogra- phy (CAS and WRIGHT 1987: p. 96). Matrix: The solid matter in which a fossil or crystal is embedded. Also, a binding substance (e.g. cement in concrete). Megabreccia: Coarse, heterolithic breccia deposits formed during caldera collapse, which contain fragments which are generally greater than one metre in diameter (LIPMAN 1976). Megabreccia fragments may be so large that individual fragments may not be readily recognizable on the scale of an outcrop. Mesobreccia: Heterolithic breccia deposits formed during caldera collapse which contain fragments that are generally less than 1 metre in diameter (LIPMAN 1976). Moat sediments: A general term for sedimentary deposits that occur between the tographic walls and the resurgent central cores of the calderas. In felsic caldera systems, moat sediments are commonly intruded by, and associated with, lava domes. Monogenetic volcano: A volcano that erupts only once (WALKER 2000: p. 283). Mudflow: A flowage of water-saturated earth material possessing a high degree of fluidity during movement. A less-saturated flowing mass is often called a debris flow. A flowing mixture composed of water and mud (clay- and silt-sized sediments). The term should be used exclusively for mud-dominated mass flows, and should not be used as a substitute for the term “lahar” (RODOLFO 2000: pp. 973–974). Mudflows are common in both volcanic and non-volcanic environments. Muddy streamflow: A flow that essentially transports sediment as normal streams do, with fine-grained sediments in suspension and coarse- grained sediment moving piecemeal along the bed as bedload (V ALLANCE 2000: p. 602). Nested caldera: A type of caldera which is found within a larger, older caldera structure. Nuees Ardentes: A French term applied to a highly heated mass of gas-charged ash which is expelled with explosive force and moves hurri- cane speed down the mountainside. The term used for a “glowing avalanche” resulting from a small-volume block and ash flow produced by the collapse of an actively growing lava dome (LACROIX 1904). In recent years, the term has unfortunately been more widely used as a synonym for “ignimbrite”. Its use should be restricted to the original definition of LaCroix (CAS and WRIGHT 1987: p. 225). Outwash: Sediments deposited by glacial meltwater beyond the active glacial ice. Outwash sediments are commonly characterized by poor- ly bedded gravels interlayered with well-bedded (and commonly cross-bedded) sands. Pahoehoe lava: A Hawaiian term to describe lava flows with smooth, continuous surfaces (KILBURN 2000: p. 291). Pahoehoe flows may have a variety of surfaces described as smooth, ropy (characterized by rope-like, commonly braided flow folds on the lava flow’s surface), or shelly (vesicular and cavernous; CAS and WRIGHT 1987: pp. 66–67). Pahoehoe toes and lobes form when largely degassed mafic magma issues from tubes relatively far from the erupting vent. Palagonite: Conspicuous yellowish alteration product of hydrated basaltic glass, mainly composed of clay minerals (SMELLIE 2000: p. 403). Pali: Hawaiian word for steep hills or cliffs. Particle cohesion: The sticking together of particles due to either the presence of water (at low temperatures) or to particle plasticity (at high temperatures) (WILSON and HOUGHTON 2000: p. 545). Pele: The mythological Polenesian goddess of volcanoes. In Hawaii, this temperamental goddess makes her home in Kilauea’s fiery vent, Halemaumau (SIGURDSSON and LOPES-GAUTIER 2000: p. 1297). Pele’s hair: A type of achnelith composed of thin, hair-like strands of volcanic glass. These thin, cylindrical strands of volcanic glass are com- monly golden in color, have diameters between 1-500 mm in diameter, and may be up to 1 metre in length. They are formed from stretched magma droplets emitted into the atmosphere during fire fountaining and strombolian eruptions (VERGNIOLLE and MANGAN 2000: p. 447). Pele’s tears: A type of achnelith composed of small droplets of shiny black volcanic glass that have been ballistically molded and quenched during flight into spherical, dumbbell, or tadpole shapes. These droplets generally range from a few millimetres to a few centimetres in size, are generally dense, but locally may be quite vesicular (VERGNIOLLE and MANGAN 2000: p. 447). 212Pelean eruption: A type of volcanic eruption characterized by a ground hugging glowing avalanche (pyroclastic flow) resulting from a mix- ture of hot volcanic gases, ash, and incandescent lava fragments. Pelean eruptions may occur when pyroclasts are blown out of a central volcanic vent and then collapse onto the earth’s surface to form a pyroclastic flow (TILLING 1985). Pelean eruptions may also occur as a result of the explosive disintegration of a lava dome (as was the case for the lava dome on Mt Pelee, Martinique in 1902). Peperite: A genetic term for a rock formed by in-situ disintigration and mixing of molten magma or lava with wet, poorly consolidated sed- iment (BATIZA and WHITE 2000: p. 361). A breccia-like deposit formed from the extrusive or intrusive mixture of lava or magma with wet sediment (SCHMIDT and SCHMINKE 2000: p. 383). Perlite: Hydrated obsidian, generally light grey in color, that is commonly characterized by rounded, onion-skin-like fractures (perlitic cracks). Apache’s tears are unhydrated clumps of fresh obsidian that are commonly found within regions containing perlite. Phreatic eruption: Phreatic eruptions are steam-driven explosions that occur when water beneath the ground or on the surface is heated by magma, lava, hot rocks, or new volcanic deposits (for example, tephra and pyroclastic-flow deposits). The intense heat of such material (as high as 1,170 °C for basaltic lava) may cause water to boil and flash to steam, thereby generating an explosion of steam, water, ash, blocks, and bombs. A steam eruption, commonly associated with water, mud, and other earth materials, that is caused when ground water, heated by a magma, flashes (and explosively expands) into steam (HARRIS 2000: p. 1301). Phreatic eruptions expel no juvenile (magmat- ic) material, and are commonly the precursor to magmatic eruptive activity. Phreatomagmatic eruption: A type of explosive volcanic eruption that occurs when water (ground water or surface water) comes in contact with hot magma. The quenching of the magma by the water causes the magma to violently fragment into juvenile (cognate) particles that are bounded by fracture surfaces and by rounded walls of broken vesicles. Due to the moisture present, accretionary lapilli are also com- mon in volcanic deposits resulting from phreatomagmatic eruptions (WILLIAMS and MCBIRNEY 1979: pp. 247–248). Piecemeal caldera: A type of caldera characterized by an internal structure composed of several individual fault-bounded blocks (LIPMAN 1997). Piecemeal calderas may result from non-uniform subsidence of a caldera formed from a single eruption, or may be the result of subsidence following a series of large eruptions (multicyclic; LIPMAN 1997; LIPMAN 2000: pp. 655–656). Pillow breccia: A mixture of coarse, typically glassy fragments and broken to whole pieces of pillow lava formed from the shattering of pil- low lava crusts (BATIZA and WHITE 2000: p. 361). Pillow breccias commonly form in areas where pillow lavas are not strong enough to maintain their competence along steep submarine slopes or scarps. Pillow lava: A type of submarine lava flow consisting of interconnected, elongated lava tubes. Cross-sections of individual lava tubes resem- ble pillows with convex upper surfaces and flat or concave lower surfaces (SCHMIDT and SCHMINCKE 2000: p. 383). Both radial and con- centric cooling fractures may be present along the margins of individual pillows, and these fractures are brought on by thermal contrac- tion during cooling. Growth of the pillow tubes takes place as the outer, commonly striated outer glassy surface of the pillow tube frac- tures, and a new tube “buds” from the fracture in a manner similar to the way that toothpaste is squeezed out of a tube. Pipe: A vertical conduit through the Earth’s crust below a volcano, through which magmatic materials have passed. Pillow interspaces are commonly filled with volcanic breccia and fragments of older rock. Pit crater: Pit craters are circular-shaped craters formed by the sinking or collapse of the ground. Fissures may erupt from the walls or base of a pit crater, but pit craters are not constructional features built by eruptions of lava or tephra. Pit craters may also partially fill with lava to form a lava lake. They are common along rift zones of shield volcanoes; for example, Mauna Loa and Kilauea volcanoes in Hawaii. No one has observed the formation of a large pit crater, but they are thought to form as a consequence of the removal of support by withdraw- al of underlying magma. Plate (piston)-type caldera: A type of caldera in which the caldera floor subsides more or less evenly as one coherent block. Plate-(piston)- type calderas are believed to result from single, large volume pyroclastic eruptions from relatively shallow depth (hypabyssal) magma chambers. Plinian eruption: Plinian eruptions are large explosive events that form enormous dark columns of tephra and gas high into the stratosphere (>11 km). Such eruptions are named for Pliny the Younger, who carefully described the disastrous eruption of Vesuvius in 79 A.D. This eruption generated a huge column of tephra into the sky, pyroclastic flows and surges, and extensive ash fall. Many thousands of people evacuated areas around the volcano, but about 2,000 were killed, including Pliny the Older. Plug: Solidified lava that fills the conduit of a volcano. It is usually more resistant to erosion than the material making up the surrounding cone, and may remain standing as a solitary pinnacle when the rest of the original structure has eroded away. Plug dome: The steep-sided, rounded mound formed when viscous lava wells up into a crater and is too stiff to flow away. It piles up as a dome-shaped mass, often completely filling the vent from which it emerged. Pluton: A body of rock which has formed beneath the earth from crystallization and consolidation from a magma (AGI 1976: p. 334). Plutons may be considered extinct magma chambers (MARSH 2000: p. 191). Large plutons (>40 square miles in area) are called “batholiths”. Polygenetic: Originating in various ways or from various sources. Pumice: Pumice is a light, porous volcanic rock that forms during explosive eruptions. It resembles a sponge because it consists of a network of gas bubbles frozen amidst fragile volcanic glass and minerals. All types of magma (basalt, andesite, dacite, and rhyolite) will form pumice. Pumice is similar to the liquid foam generated when a bottle of pressurized soda is opened — the opening depressurizes the soda and enables dissolved carbon dioxide gas to escape or erupt through the opening. During an explosive eruption, volcanic gases dissolved in the liquid portion of magma also expand rapidly to create a foam or froth; in the case of pumice, the liquid part of the froth quickly solidifies to glass around the glass bubbles. Solidified fragments of quenched, highly vesicular (>60%) silicic magma or lava (CASHMAN et al. 2000: p. 421). The highly vesicular nature of pumice results from large volumes of gas rapidly expanding within a rapidly cooling magma. The low density of pumice commonly permits it to float on water for extended periods of time. Hot pumice, however, has been shown experimentally to sink rapidly upon interacting with water (WHITHAM and SPARKS 1986 ). Pyroclastic: Refers to processes resulting from the explosive fragmentation of a magma or lava. May also be used to describe the deposits formed by explosive volcanic activity and directly deposited by transport processes resulting directly from this activity (CAS and WRIGHT 1987: p. 8). Pyroclastic is a Greek term which means “fire-broken” (HARRIS 2000: p. 1301). Pyroclastic breccia: This is produced by volcanic explosion and includes vulcanian breccia, pyroclastic flow breccia, and hydrovolcanic brec- cia (FISHER 1960). Pyroclastic density current: A gravity controlled, laterally moving mixture of pyroclasts and gas (WILSON and HOUGHTON 2000: p. 545) Pyroclastic fall: The “rain-out” of pyroclasts through the atmosphere from an eruption jet or eruption plume during an explosive volcanic eruption (WILSON and HOUGHTON 2000: p. 545; HOUGHTON et al 2000: p. 555). 213Pyroclastic fall deposit: Volcaniclastic (pyroclastic) deposits formed from the rain-out of clasts through the atmosphere from an eruption jet and/or plume during an explosive eruption (HOUGHTON et al. 2000: p. 555). Fall deposits typically exhibit mantle bedding, are well sort- ed, and commonly show well-developed planar stratification (CAS and WRIGHT 1987: pp. 95–96). Pyroclastic flow: A dense, hot, dry, high particle concentration mixture of gas and hot rock fragments (ash, pumice, blocks, etc.) that travels along the ground surface, typically at high velocity (generally on the order of hundreds of feet or metres per second; HARRIS 2000: p. 1301) away from a volcano. The high speeds of pyroclastic flows are possible because they flow over a thin layer of hot, commonly expanding and escaping gases. Most of the material within a pyroclastic flow is contained within concentrated particle dispersion locat- ed at the flow’s base (WILSON and HOUGHTON 2000: p. 545). Pyroclastic flow deposit: Pyroclastic (volcaniclastic) deposits that are left by pyroclastic flows (CAS and WRIGHT 1987: p. 96). The deposits are usually topographically controlled (infilling valleys and topographic depressions), massive, and poorly sorted. Depending upon their thickness and heat retention, pyroclastic flow deposits may coalesce into welded tuffs. Pumice-rich pyroclastic flow deposits are often called “ignimbrites”. Pyroclastic surge: A type of turbulent, low density (low particle concentration) pyroclastic cloud or pyroclastic density current. Being more dilute than pyroclastic flows, surges can sweep over ridges, hills, and other topographic boundaries. Two kinds of surges are known: wet surges have temperatures <100 °C and contain steam that condenses into water droplets that surge along the ground surface with gas and pyroclasts; and dry surges, which have temperatures >100 °C, and form by either hydrovolcanic eruptions with low water/magma ratios, or by magmatic erup- tions driven solely by expanding magmatic gases (V ALENTINE and FISHER 2000: p. 571). Pyroclastic surge deposit: Pyroclastic deposits that are left by pyroclastic surges. These deposits mantle topographic features but also gen- erally thicken within topographic depressions. These deposits are generally well-sorted, and are enriched in crystals and lithic frag- ments relatively to pyroclastic flow deposits. Surge deposits commonly exhibit unidirectional sedimentary bedforms, including low angle cross-bedding, dune forms, climbing dune forms, pinch and swell structures, and chute and pool structures (CAS and WRIGHT 1987: p. 98). Quenching: The rapid cooling of magma to form glass (BATIZA and WHITE 2000: p. 361). Fuel-coolant interactions commonly lead to quench- ing. Abrupt quenching may cause a rapid volume decrease which leads to fragmentation of the glass (cooling-contraction granulation). Renewed volcanism state: Refers to a state in the evolution of a typical Hawaiian volcano during which — after a long period of quiescence — lava and tephra erupt intermittently. Erosion and reef building continue. Repose: The interval of time between volcanic eruptions. Reticulite: Reticulite is basaltic pumice in which nearly all cell walls of gas bubbles have burst, leaving a honeycomb-like structure. Even though it is less dense than pumice, reticulite does not float in water because of the open network of bubbles. The delicate glass threads between the bubbles are so fragile that reticulite was first called “thread-lace scoria” by the great American mineralogist, James Dana. An exceptionally porous type of scoria containing porosities ranging from 95-99% (VERGNIOLLE and MANGAN 2000: p. 447; MCPHIE et al. 1993: p. 27). Commonly referred to as “thread-lace” scoria, reticulite is made up of a honeycomb-like network of thin glass fibers. Resurgent dome: The central highland in many large calderas formed by gradual upwarping of the caldera floor after caldera collapse as a result of renewed magma intrusion. Reynolds number: A dimensionless quantity characterizing the relative importance of inertial or momentum related forces to viscous forces in fluid flow. In magmas viscous forces usually dominate, resulting in a Reynolds number that is less than 1 (CARRIGAN 2000: p. 219) Ring fracture/Ring fault: The arcuate bounding faults upon which caldera (cauldron) subsidence takes place. Ring fractures (faults) define the structural limits of calderas. Most observed ring faults are nearly vertical or dip steeply inward (toward the center of the caldera), and this is thought to be a result of doming of the caldera structure following its initial formation (LIPMAN 2000: pp. 649–650). Ring plain: Region surrounding a volcano beyond lower topographic flanks, over which tephra and mass-wasting products are radially dis- tributed. Ropy pahoehoe: A type of pahoehoe lava characterized by flexible crusts that are bent into tight folds as lava flows. These tight folds form lava surfaces that appear to be made up of a series of braided ropes (KILBURN 2000: p. 295). Scoria: Scoria is a vesicular (bubbly) glassy lava rock of basaltic to andesitic composition ejected from a vent during explosive eruption. The bubbly nature of scoria is due to the escape of volcanic gases during eruption. Scoria is typically dark grey to black in colour, mostly due to its high iron content. The surface of some scoria may have a blue iridescent colour; oxidation may lead to a deep reddish-brown colour. Solidified fragments of quenched, highly vesicular (>60%) mafic magma or lava (CASHMAN et al. 2000: p. 421). The highly vesicular nature of scoria results from rapid cooling of gas-rich lava. Scoria (cinder) cone: Small volcanic landforms formed from focused (single-vent) subaerial strombolian eruptions of basalt or basaltic- andesite magma. These features have an inverted cone-shaped profile and are generally circular in plan, although elongate scoria cones can be formed from multiple-vent volcanic eruptions (CAS and WRIGHT 1987: pp. 371–372). Seamount: A submarine volcano. Sector collapse: A destructive volcanic process during the growth history of a volcano. Debris avalanche deposits are the products of sector collapses (UI et al. 2000: p. 617). Shearing: The motion of surfaces sliding past one another. Shelly pahoehoe: A type of pahoehoe lava characterized by highly vesicular, extremely fragile crusts that form over hollow lava blisters. The surfaces of these blisters break easily when stepped upon, giving the impression of walking on eggshells (KILBURN 2000: p. 295). Shield volcano: A broad, low-relief volcano constructed by flows of relatively fluid lava (e.g. basalt: SPUDIS 2000: p. 698). Flank slopes on shield volcanoes are typically < 5° (ZIMBELMAN 2000: p. 771). Sill: A tabular body of intrusive igneous rock, parallel to the layering of the rocks into which it intrudes. Sinter: A type of fragile, commonly white or grey rock formed by precipitation of silica from cooling hydrothermal solutions at or near a hydrothermal vent. Precipitation of siliceous sinter (often with associated sulphide minerals and precious metals) commonly occurs in neutral and acid hydrothermal systems under the influence of biogenic agents such as algae and bacteria (CAS and WRIGHT 1987: p. 316). Skylight: An opening formed by a collapse in the roof of a lava tube. Slabby pahoehoe: A type of pahoehoe lava with a surface composed of slabs of broken lava crust that are up to metres across and up to sev- eral centimetres thick (KILBURN 2000: p. 295). Slug flow: One of four two-phase flow regimes, in which large gas pockets,which are almost as large as the eruption conduit rise through magma (JAUPART 2000: p. 237). 214Solfatara: A type of steam vent or dry fumarole that is characterized by quiet discharge (<20 m/s), and that precipitates a significant amount of sulphur (HOCHSTEIN and BROWNE 2000: pp. 850-851). Spall fragments: Formed by shattering of the brittle, quenched, crust of a lava by a combination of thermal contraction and flexure at the margins of a still-flowing lava (BATIZA and WHITE 2000: p. 361) Spatter bomb: A glassy pyroclast greater than 64 mm in diameter that takes on a fluidal shape by the force of ejection (VERGNIOLLE and MANGAN 2000: p. 447). Spatter cone: A low, steep-sided cone of spatter built up on a fissure or vent. It is usually of basaltic material. Spatter rampart: A ridge of congealed pyroclastic material (usually basaltic) built up on a fissure or vent. Spherulite: Typically rounded, radiating arrays of crystal fibers produced by the high temperature devitrification of volcanic glass. In felsic rocks, the crystal fibers are generally composed of alkali feldspar and a silica polymorph (either quartz or cristobalite), whereas in mafic rocks the fibers commonly consist of plagioclase and/or pyroxene. Spherulites typically have diameters of 0.1-2.0 cm, but can be much larger (commonly up to 20 cm). Isolated spherulites are generally spherical, but adjacent spherulites may impinge upon one another to produce long chains that are often aligned with flow foliation (MCPHIE et al. 1993: pp. 24–25). Spindle bomb: Volcanic bombs are masses of lava greater than 64 mm diameter ejected from a vent. Bombs are viscous when expelled and deform during flight to assume a characteristic shape on impact. Spines: Horn-like projections formed upon a lava dome. Steady-state or equilibrium profile: Shape of the edifice (cone) once an active volcano has become well established — follows the initial cone building, precedes long-term erosional degradation, and represents a balance between construction through mass addition (eruption) and degradation through erosion. Stratovolcano: A generally steep sided volcano composed of alternating layers of lava flows, pyroclastic deposits, and commonly, volcani- clastic sedimentary deposits (WALKER 2000: p. 283). Stratovolcanoes commonly have increasing slopes toward their summits since they generally have mainly lava flows and sedimentary deposits near their base and pyroclastic (tephra) deposits near their summits. Stratovolcanoes also called as “composite volcano”. Stony rhyolite: Very finely crystalline rhyolite lava (CAS and WRIGHT 1987: p. 84). Strombolian eruption: Volcanic eruptions of basaltic magma, slightly more violent than Hawaiian eruptions, that produce large amounts of scoria and ash around a central vent to form a cone. Strombolian eruptions are typically pulsating and have periods of several seconds (WOLF and SUMNER 2000: p. 321). The deposits consist of lava spatter, vesicular bombs, scoria lapilli, and mafic ash (VESPERMANN and SCHMINCKE 2000: p. 683). It named after Stromboli, an Italian volcano. Superheated liquid: A metastable thermodynamic state of a liquid resulting from rapid heating to a temperature well above the boiling point (MORRISSEY et al. 2000: p. 431). Surge: A ring-shaped cloud of gas and suspended solid debris that moves radially outward at high velocity as a density flow from the base of a vertical eruption column accompanying a volcanic eruption or crater formation. A pyroclastic surge is a turbulent cloud of gas and rock fragments that flows across the ground. A pyroclastic surge is more dilute than a pyroclastic flow. Surges are not constrained by topography but can move over obstacles such as ridges and hills. Surtseyan eruptions: Hydrovolcanic eruptions dominated by jets of wet tephra (scoria and ash) that result in the formation of tuff cones. The term “surtseyan” is generally used for volcanoes erupting through seawater. Named after Surtsey, a volcano which emerged from the sea off the coast of Iceland in 1963 (VESPERMANN and SCHMINCKE 2000: p. 683). Talus: A slope formed a the base of a steeper slope, made of fallen and disintegrated materials. Tephra: A general term used by volcanologists to describe all fragmental volcanic ejecta produced during explosive volcanic eruptions (DEHN and MCNUTT 2000: p. 1271). This includes ash (<2 mm diameter fragments), lapilli (2–64 mm diameter fragments and fragements greater than 64 mm in diameter known as bombs (semi-solid or plastic ejecta) or bombs (solid ejecta) (TILLING et al. 1987). Tephra (finger) jet: parcels of bombs, wet tephra, and gas that follow spreading parabolic trajectories and are ejected by discrete, relatively low frequency shallow explosions (SMELLIE 2000: p. 403) Tephrochronology: The collection, preparation, petrographic description, and approximate dating of tephra. Topographic inversion: Process whereby through time valleys become ridges and vice versa — can occur on volcanoes as volcanogenic prod- ucts such as lavas are channeled down valleys, focusing subsequent erosion along their edges. Trap-door caldera: A type of caldera formed when one part of the caldera floor subsides to a greater depth than the other side of the caldera floor. In general, trap-door calderas have a partial ring fracture (associated with the side of greatest caldera collapse) and a hinge area (associated with the side of least collapse). Trap-door calderas may represent either calderas that have undergone incomplete collapse, or calderas formed from eruptions from shallow asymmetrical magma chambers (LIPMAN 1997; LIPMAN 2000: p. 654). Tree mold: Fluid basaltic lava may preserve the shapes of trees and other objects by solidifying around them. Tree molds are formed when lava surrounds a tree, chills against it, and then drains away. The standing structure left behind is often called a lava tree. Tree trunks engulfed and incinerated by lava leave cylindrical hollows, or tree molds, where lava solidified against them; tree molds often preserve the original surface texture of the tree. Tree molds are found within standing lava trees and on the surfaces of lava flows. They are com- mon in pahoehoe flows and occasionally found in `a`a flows. Tremor: A continuous vibration of the ground around active volcanoes (VERGNIOLLE and MANGAN 2000: p. 447). Tremors defined on seis- mographs may have either a regular sine-wave appearance (harmonic tremor) or an irregular, pulsating appearance (spasmodic tremor) (MCNUTT 2000: p. 1015). Tu f f : A lithified volcaniclastic rock composed primarily of ash, with up to minor volumes of lapilli and/or blocks and bombs (FISHER 1966). Originally used as a non-genetic rock name, common use today typically implies (incorrectly) that the tephra comprising the rock was deposited while hot. Similar deposits that have no indication of being hot while deposited are commonly referred to as “tuffaceous” (MCPHIE et al. 1993: p. 8). Tuff cone: A type of hydroclastic volcano that is generally higher than (generally >50 m high), and has steeper external flanks (commonly >25°) than tuff rings or maars (VESPERMANN and SCHMINCKE 2000: p. 684). Craters within tuff cones are generally higher in elevation than the adjacent land surface. Tuff cones are made up primarily of juvenile clasts deposited from lateral surges, airfall, and associated volcaniclastic remobilization processes. Tuff ring: A type of hydroclastic volcano, generally <50 m high, defined by craters with low depth/width ratios that sit at or above the eleva- 215tion of the adjacent land surface. The rims around tuff rings are composed of juvenile and accidental clasts and are deposited in beds with dips <25° (VESPERMANN and SCHMINCKE 2000: p. 684). Tumescence: The doming or uprising of a volcano commonly due to inflation of a shallow magma chamber. Regional tumescence common- ly occurs prior to a major pyroclastic eruption, but may also occur following an eruption as less volatile magma is emplaced into the shal- low crust (SMITH and BAILEY 1968). Tumulus: The surfaces of pahoehoe flows on flat or gentle slopes often exhibit elliptical, domed structures called tumuli. A tumulus is cre- ated when the upward pressure of slow-moving molten lava within a flow swells or pushes the overlying crust upward. Since the solid crust is brittle, it usually breaks to accommodate the “inflating” core of the flow. Such fractures generally extend along the length of a tumulus, and are frequently accompanied by smaller irregular cracks down the sides. Lava commonly squeezes out through these frac- tures, and sometimes drains from the tumulus to leave a hollow shell. Turbidite: Sediment or rock deposited from a gravity-driven flow of suspended sediment in water (turbidity current), characterized by well- developed sedimentary structures arranged in a regular sequence (SMELLIE 2000: p. 404) Turbulent flow: flow regime where inertial effects dominate, such that the flow is well mixed by small-scale eddies (JAUPART 2000: p. 237) Tu ya : A flat-topped, steep-sided volcano that erupted into a lake thawed into a glacier by volcanic heat (SMELLIE 2000: p. 403). Tuya com- monly referred to as a “table mountain”. Unconformity: A surface of erosion that separates younger strata from older rocks (AGI 1976: p. 448). U-shaped channels: These channels in base surge deposits, described by several authors, are symmetrical in cross section with curving bot- toms that clearly cut underlying layers. Most range from about 0.3 m to 7 m across and are a few centimetres to 3 m deep, but unusually large channels (30 m across 20 m deep) are also reported. The curving bottoms are best described as U-shaped, not parabolic curves, even though some are very broad in cross section. Beds reflect the shape of the channels, but the curvature of individual beds decreases upward, and the final fill extends uniformly across the channel and is conformable with the sequence outside the channel. Thus, beds thicken toward the centers of channels and therefore do not resemble draped fallout layers. Vapor film: A layer of vapour formed at the interface of water and hot liquid or solid body (MORRISSEY et al. 2000: p. 431). Variolite: A spherulite-like radiating aggregate composed of feathery, needle-like crystals of plagioclase and pyroxene that occur in mafic volcanic rocks (typically basalt). Variolites may result from devitrification, but are commonly believed to be formed in subaqueous rocks by quench-induced crystallization (CAS and WRIGHT 1987: p. 420). VEI: The Volcanic Explosivity Index, or VEI, was proposed in 1982 as a way to describe the relative size or magnitude of explosive volcanic eruptions. It is a 0-to-8 index of increasing explosivity. Each increase in number represents an increase around a factor of ten. The VEI uses several factors to assign a number, including volume of erupted pyroclastic material (for example, ashfall, pyroclastic flows, and other ejecta), height of eruption column, duration in hours, and qualitative descriptive terms. Vent: A surface opening through which volcanogenic materials are erupted (DA VIDSON and DESILVA 2000: p. 663). Typically thought of as a hole in a planet from which volcanic products (magma, ash, etc.) are erupted (SPUDIS 2000: p. 697). Vesicle: A frozen bubble in a volcanic rock. Vesicles are formed when magma crystallizes around a gas bubble (SPUDIS 2000: p. 697). Vesicular tuff: Tuffs containing millimetre to centimetre-sized, irregular to round vesicles which are interpreted to form during trapping of air or vapor in wet ash deposits (VESPERMANN and SCHMINCKE 2000: p. 683). Vesiculation: Nucleation and growth of gas bubbles in a magma (CASHMAN et al. 2000: p. 421) Vesuvian eruption: Commonly used as a synonym for a “Plinean” eruption (e.g. TILLING 1985), but also used to describe basaltic eruptions which involve long-sustained gas streaming with little ash being released (as in the 1906 eruption of Vesuvius; CAS and WRIGHT 1987: p. 130). Viscosity: A measurement of the ratio of shear stress to the rate of shear strain in a fluid (WILLIAMS and MCBIRNEY 1979: p. 20). In common language, how easily a fluid flows. Considered the most important physical property of a magma because it largely determines eruptive style as well as volcano morphology. Magma viscosity generally increases as the silica content of the magma increases (due to silica poly- merization) and as the temperature of the magma decreases. Magma viscosity may also be affected by the presence of trace elements (e.g. Ti) or volatiles (e.g. H 2 O, CO 2 , SO 2 , etc.). In general, common magmas increase in viscosity in the following order: komatiite, basalt, andesite, dacite, rhyodacite, rhyolite. Vog: Smog of volcanic origin, composed of volcanic ash and gases Volcanic field: These are clusters of small volcanoes which erupt only once. They usually have a cinder cone and lava flows. A typical field may contain 10 to 100 volcanoes. One of the largest is the Michoacan–Guanajuato field in Mexico which contains 1000 volcanoes. Volcanic bomb: Juvenile fragments of semi-solid or plastic magma ejected during a volcanic eruption. Based on their shapes after they hit the ground and cool, bombs are given various textural names including breadcrust bombs, cow-dung (cow pie) bombs, spindle bombs (fusiform bombs) and ribbon bombs. Type of bombs: Rotational bomb, Tear-shaped bomb, Spheroidal bomb, Spindle-shaped bomb, Pancake-shaped bomb, Slag bomb, Pumiceous bomb, Ribbon bomb, Cored bomb, Olivine bomb, Breadcrust bomb, Cow pat bomb Volcanic breccia: The term volcanic breccia is used as a general term applying to all coarse-grained rocks composed of angular volcanic frag- ments (FISHER 1960). Volcanic complex: A persistent volcanic vent area that has built a complex combination of volcanic landforms. Volcanic cone: A mound of loose material that was ejected ballistically. Volcanic cycle: A general term used to describe a period of increased volcanic activity. Volcanic field: A region comprising a large number of volcanic edifices. Volcanic fields are usually associated with basaltic volcanism, and commonly comprise a number of small, monogenetic volcanoes (e.g. cinder cones, maars, tuff cones, tuff rings, small shield volcanoes, lava domes). Fields may form in linear trends associated with tectonic structures (such as faults), on the flanks of larger composite or shield volcanoes, or within calderas (CONNOR and CONWAY 2000: p. 331). Volcanic magnitude: Volcanic magnitude is the total mass of erupted material. The scale spans from 0 to 9 magnitude = log10 (erupted mass kg) –7. The magnitude will be similar to VEI which also measures mass erupted. Volcanic neck: A massive pillar of rock more resistant to erosion than the lavas and pyroclastic rocks of a volcanic cone. Volcaniclastic: A non-genetic term used to describe any fragmental aggregate of volcanic parentage (CAS and WRIGHT 1987: p. 8). Rocks formed by the fragmentation of volcanic materials (either magma or volcanic rocks) irrespective of the method of fragmentation. Pyroclastic rocks and epiclastic rocks are both considered to be “volcaniclastic”. 216Volcaniclastic facies: These facieses are defined by distance from source, type of transporting agent, environment of deposition, and in some cases, by composition. First-order volcaniclastic facies are generally defined by position of the rock body relative to source within non- marine or marine environments, e.g. proximal, medial and distal facies. Vulcan: Roman god of fire and the forge after whom volcanoes are named. Vulcanian eruption: A vulcanian eruption is a type of explosive eruption that ejects new lava fragments that do not take on a rounded shape during their flight through the air. This may be because the lava is too viscous or already solidified. These moderate-sized explosive erup- tions commonly eject a large proportion of volcanic ash and also breadcrust bombs and blocks. Andesitic and dacitic magmas are most often associated with vulcanian eruptions, because their high viscosity (resistance to flow) makes it difficult for the dissolved volcanic gases to escape except under extreme pressure, which leads to explosive behavior. An explosive volcanic eruption generally expelling less than 1km3 of material, but with an eruption column that may reach heights of up to 10-20 km (NAKADA 2000: p. 945). These eruptions last on the order of seconds to minutes (MORRISSEY and MASTIN 2000: p. 463). Welding: The sintering together of hot, glassy fragments, irrespective of shape and size, by compactional lithostatic load (CAS and WRIGHT 1987: p. 165). Postdepositional compaction of a hot pyroclastic deposit under its own weight. Note that agglutination, coalescence, and postdepositional welding are part of a process continuum (WOLFF and SUMNER 2000: p. 321). Welded tuff: A hard pyroclastic rock compacted by internal heat and pressure from overlying pyroclastic deposits. Wet/dry eruptions: Depending on the quantity of available heat energy, when water comes into contact with magma it may become a liquid- rich fluid or a vapour-rich fluid, resulting in a dry eruption or a wet eruption respectively (MORRISSEY et al. 2000: p. 431). Xenocrysts: A crystal that resembles a phenocryst in igneous rock, but is a foreign to the body of rock in which it occurs. Xenoliths: A foreign inclusion in an igneous rock. References AGI 1976: Dictionary of Geological Terms. — Garden City, New York, American Geological Institute, Anchor Press, 472 p. BATIZA, R. and WHITE, J. D. L. 2000: Submarine lavas and hyaloclastite. — In: SIGURDSSON, H., HOUGHTON, B. F., MCNUTT, S. R., RYMER, H. and STIX, J. (Eds): Encyclopedia of Volcanoes. 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San Diego, Academic Press, pp. 771–783. 218Aa (Ch5) 107 Aa lava (Ch2) 39 Accessory (Ch1) 8, 10 Accidental (Ch1) 8, 10, 20 Accretionary (Ch3) 63, 64 Agglutinates (Ch3, Ch5) 68, 108 Apron (Ch4) 96 Armoured (Ch3) 63 Autobreccia (Ch1) 14 Autoclastic (Ch1, Ch3) 13, 14, 15, 21, 57, 65 Ballistic clasts (Ch3) 69 Base surge (Ch3, Ch4, Ch8) 63, 68, 88, 195 Bedding (Ch3) 63, 67 Block (Ch7) 162, 165, 166, 167, 168, 169 Block-and-ash flows (Ch3, Ch6) 68, 69, 136, 137 Brittle fragmentation (Ch8) 191 Bubble coalescence (Ch3) 64 Bubbles (Ch3) 58, 61, 64 Bulking (Ch7) 162, 163 Buoyancy (Ch5) 105 Caldera (Ch6) 134, 135, 138, 139, 140, 141 Carbonatites (Ch5) 111 Cavity (Ch8) 199 Central volcano (Ch5) 105, 109 Clast morphology (Ch1) 21 Clastic (Ch4) 87, 89, 92 Clastic dykes (Ch3) 63, 70 Cock’s tail jets (Ch5) 118 Cognate (Ch1) 10 Coherent (Ch4) 87, 89 Cohesion (Ch3) 66, 67, 68 Columnar joint (Ch2) 39, 40 Componentry (Ch1, Ch4) 8, 9, 11, 21, 89 Composite clasts (Ch1) 8, 21 Composite volcano (Ch6) 133, 134, 135, 135, 139 Conductivity (Ch8) 190, 191 Conduit (Ch5) 103, 105, 109, 111, 117 Continental (Ch4) 88, 89 Country rock (Ch5) 103, 104, 105, 110, 111, 112, 119 Crater lake (Ch5) 106, 111, 114, 118 Cross-bedding (Ch1) 11 Cryptodomes (Ch8) 189, 190 Crystalisation (Ch3) 61 Crystals (Ch1) 9, 10 Cyclic (Ch6) 139 Debris avalanches (Ch3, Ch7) 67, 70, 159, 163, 165, 166, 167, 168, 171 Debris flows (Ch3, Ch5, Ch7) 67, 113, 114, 159, 161, 162, 163, 164, 168, 169, 170 Decompression (Ch3) 57, 61 Degassing (Ch4) 94 Density (Ch5) 105, 115 Density currents (Ch5, Ch8) 103, 105, 113, 117, 118, 191, 192, 193, 195 Density separation (Ch4) 89 Devitrification (Ch3) 65, 66 Diatreme (Ch5) 105, 106, 110, 111, 112, 113, 114, 115, 116, 117, 119 Direct blast (Ch3) 58 Distal (Ch4) 87, 88, 91, 94 Dry surges (Ch3) 68 Dune-bedded (Ch1) 10 Dyke (Ch2) 37 En masse freezing (Ch3, Ch6) 68, 137 Epiclastic (Ch1) 13, 15, 20 Erosion (Ch7) 159, 163, 168, 169, 170, 171 Erosional (Ch4) 87, 93 Eruption cloud (Ch3) 58, 63, 67, 68 Eruption-fed (Ch8) 193 Explosion locus (Ch5) 110 Facies analysis (Ch4) 87, 89, 93, 94 Feeder dyke (Ch5) 111, 116 Feeding (Ch5) 105, 111 Flow regimes (Ch2, Ch3) 11, 68, 69 Flow steadiness (Ch3) 68 Flow/current stability (Ch3) 66 219Flow-banded (Ch3) 65, 66 Fountaining (Ch5) 109, 117 Fragmentation (Ch1, Ch3) 7, 8, 9, 10, 13, 15, 57, 58, 59, 60, 61, 62, 63, 64, 65, 66 Fuel-coolant interaction (Ch3) 59 Gas thrust (Ch3) 58, 59 Genetic (Ch1) 7, 8, 15, 18, 19, 20, 21 Glass (Ch3) 59, 61, 62, 63, 64, 65, 66 Grading (Ch1) 10, 11 Grain avalanches (Ch3) 67 Grain flows (Ch3) 67 Granular flows (Ch3) 68 Gravity current (Ch8) 193 Gravity separation (Ch4) 89 Hard rock (Ch5) 111 Hawaiian (Ch2, Ch5) 32, 33, 34, 106, 107, 108, 109, 117 Heat capacity (Ch8) 190, 191 High grade (Ch6) 138 Hole-in-the-ground (Ch5) 105, 110 Horizontal (Ch4) 89, 91, 93, 94, 95 Hummocks (Ch7) 164, 166 Hyaloclastic (Ch1) 14, 21 Hyaloclastite (Ch2, Ch3, Ch6, Ch8) 41, 42, 65, 66, 70, 134, 187, 188, 189, 190, 192, 197, 199 Hydroclastic (Ch1) 13 Hydrothermal (Ch6) 136, 140, 141, 142 Hyperconcentrated flows (Ch7) 163, 168, 169 Hyperconcentrated mass flows (Ch5) 113 Ignimbrite (Ch4, Ch6) 88, 93, 94, 96, 138, 141, 142 Impact sags (Ch3, Ch5) 69, 113 Inter-eruptive (Ch4) 89 Intracontinental (Ch5) 104, 107 Jig-saw fit texture (Ch7) 165, 166 Jökulhlaup (Ch8) 197 Juvenile (Ch1) 8, 9, 10, 21 Kimberlites (Ch5) 111 Laccolith (Ch2) 37, 38, 40 Lacustrine (Ch8) 187, 188, 192, 197 Laharic (Ch4) 88, 93 Lahars (Ch7) 159, 161, 162, 163, 164, 165, 167, 168, 171 Laminar (Ch2) 32 Laminated (Ch5) 114 Landslide (Ch7) 164, 165, 167 Large igneous provinces (Ch2) 38 Lava dome (Ch2, Ch6, Ch8) 39, 40, 41, 42, 135, 136, 137, 189, 190 Leidenfrost effect (Ch3) 59 Levees (Ch7) 166 Limu (Ch8) 195 Lithofacies (Ch1, Ch4) 18, 19, 89, 90, 91, 92 Lithological (Ch1) 15, 18, 19 Lithophysae (Ch3) 65, 66 Lobate lava (Ch8) 188 Low grade (Ch6) 138 Maar (Ch2, Ch5) 30, 103, 105, 106, 108, 109, 110, 111, 112, 113, 114, 115, 116, 117, 119 Magma mixing (Ch3) 61 Magmatic fragmentation (Ch3) 57, 58, 59, 60, 61, 63, 64, 65 Magmatic volatiles (Ch3) 57 Mass flow (Ch3, Ch7) 66, 67, 161, 162, 163, 164, 165, 166, 167, 168 Matrix (Ch7) 164, 165, 166, 168 Mega-calderas (Ch6) 135, 139, 141 Microlite (Ch1, Ch3) 10, 62, 63, 64, 66 Mid-ocean ridges (Ch8) 187, 188 Mofetta (Ch6) 141 Molten-fuel-coolant-interaction (Ch8) 190 Monogenetic (Ch2, Ch5, Ch6) 29, 30, 34, 37, 39, 40, 103, 104, 105, 106, 110, 119, 134 Mud flow (Ch3) 67 Muddy slurry (Ch3) 59, 60 Obsidian (Ch2, Ch3) 41, 65 Overbank (Ch4) 93 Pahoehoe (Ch2, Ch5) 39, 41, 107 Palagonite (Ch3) 65 Parallel bedded (Ch1) 10 Particle concentration (Ch3) 58, 63, 66, 67, 68, 69 Pele’s hair (Ch3) 64 Peperite (Ch1, Ch5, Ch8) 21, 111, 115, 116, 117, 188, 189 Perlitic (Ch2, Ch3) 41, 66 Phreatic explosion (Ch6) 142 Phreatomagmatic fragmentation (Ch3) 57, 58, 59, 60, 61, 63 Phreato-plinian (Ch2) 33, 35 Pillow (Ch8) 188, 189, 192 Pillow lava (Ch2) 41 Plinian (Ch2, Ch5) 30, 32, 33, 34, 35, 36, 107, 109 Polygenetic (Ch2, Ch5, Ch6) 29, 105, 133, 134, 141 Premixing (Ch8) 190 Primary volcaniclastic (Ch1) 8, 13, 20, 21, 22 Progressive aggradation (Ch3, Ch6) 67, 68, 137 Proximal (Ch4) 87, 88, 91, 94 P-type (Ch2) 39 Pumice (Ch3, Ch8) 59, 61, 64, 191, 195, 196 Pyroclastic density current (Ch1, Ch3, Ch4, Ch5) 10, 11, 63, 68, 88, 89, 93, 94, 105, 113 Pyroclastic fall (Ch2, Ch3) 31, 32, 66, 67, 68 Pyroclastic flow (Ch2, Ch3, Ch6, Ch7, Ch8) 31, 32, 34, 58, 67, 68, 69, 70, 135, 136, 137, 138, 139, 161, 162, 164, 168, 191 Pyroclastic mounds (Ch5) 105 Pyroclastic rocks (Ch1) 7, 8, 14 Pyroclastic surge (Ch2, Ch3) 31, 32, 34, 58, 66, 67, 68, 69 Pyroclasts (Ch3) 57, 58, 59, 61, 63, 64, 65, 66, 67 Pyrogenic (Ch1) 10 Quench granulation (Ch8) 191 Rain falls (Ch7) 162 Recurrence (Ch5) 105, 119 Remobilisation (Ch5) 103, 113 Resedimentation (Ch1) 16, 17 Resurgence (Ch6) 134, 139, 141 Reworked (Ch3) 57, 64, 67, 70 Reworking (Ch1) 15, 17 220Ring plain (Ch6, Ch7) 134, 136, 159, 160, 161 Rock falls (Ch3) 67 Root zone (Ch5) 105, 110, 111, 112 Sandwave (Ch5) 113 Scoria (Ch3) 59, 64 Scoria cone (Ch2, Ch5) 30, 31, 34, 40, 109 Scoria-and-ash flow (Ch6) 136, 138 Seamounts (Ch8) 188 Secondary (Ch4) 87, 88, 89, 94 Sector collapse (Ch7) 165, 171 Segregation (Ch3, Ch7) 63, 163 Shield (Ch2) 29, 31, 40 Shield volcano (Ch5, Ch6) 104, 107, 133, 134 Shock waves (Ch3) 59, 61 Sideromelane (Ch1, Ch3) 9, 10, 61, 62, 64 Sill (Ch2) 37, 38, 40 Slurry (Ch8) 192 Small-volume (Ch5) 103, 104, 105, 106, 110, 115 Snow slurry lahars (Ch7) 162 Soft substrate (Ch5) 111, 112 Soft-sediment deformation (Ch3) 70 Solfatara (Ch6) 141 Sorting (Ch1) 11, 12, 21 Spall dome (Ch8) 192 Spatter cones (Ch5) 105, 106, 107, 108, 115 Spherulites (Ch3) 65, 66 Stratigraphy (Ch4) 88, 93 Strato-volcano (Ch6) 134 Stream flow (Ch7) 164, 168 Strombolian (Ch2, Ch5) 32, 33, 34, 106, 107, 108, 109, 110, 116, 117 S-type (Ch2) 39 Subglacial (Ch8) 187, 192, 197 Sublacustrine (Ch5) 105 Sub-Plinian (Ch2) 30, 33, 34, 35 Subsidence (Ch5) 114, 119 Superheated (Ch3) 59 Surge (Ch3) 58, 63, 66, 67, 68, 69 Surtla (Ch8) 195 Surtseyan (Ch2, Ch5, Ch8) 33, 36, 117, 118, 119, 192, 199 Suspension (Ch3) 58, 66, 67, 68 Suspension-dominated transport (Ch3) 66 Sustained pyroclastic flow (Ch3) 67 Syn-eruptive (Ch4, Ch5) 89, 103, 105, 106, 114 Tachylite (Ch1, Ch3) 9, 10, 62, 64 Tephra jetting (Ch5, Ch8) 118, 195 Thermohydraulic (Ch5) 111 Thermohydraulic explosion (Ch3) 59 Tindar (Ch8) 192, 199 Traction-dominated transport (Ch3) 66, 67 Trajectory (Ch3) 66, 67, 69 Transitional flows (Ch7) 161 Tuff cones (Ch5) 103, 105, 117, 118 Tuff ring (Ch2, Ch5) 30, 103, 105, 113, 114, 116, 117 Tumuli (Ch2, Ch5) 39, 41, 107 Turbidity currents (Ch3, Ch5) 67, 68, 114 Turbulent (Ch4) 89, 94 Turbulent mixing (Ch4) 89 Tuya (Ch8) 199 Uprush (Ch5) 118 Valley filling (Ch7) 162, 164, 168 Vapour film (Ch3) 59 Vault (Ch8) 199 Vent geometry (Ch5) 103 Vertical (Ch4) 89, 91, 92, 93, 94 Vesicularity (Ch4) 94 Vesiculation (Ch1, Ch3) 9, 10, 57, 58, 59, 61 Volcanic collapses (Ch7) 159, 166, 167 Volcanic sediment (Ch4) 87, 89, 92, 93 Volcanic successions (Ch1) 7 Volcanic textures (Ch1) 7 Volcaniclastic (Ch3) 57, 64, 67, 69, 70 Volcaniclastic rock (Ch1) 7, 8, 13, 14, 15, 20, 21, 22 Volcano collapse (Ch6) 133, 135 Volcanogenic (Ch1) 17, 19 Vulcanian (Ch2) 33, 34 Water to magma ratio (Ch3) 60 Welded (Ch8) 191, 193 Wet surges (Ch3) 68 Zonation (Ch4) 94 221