Volkanoloji Volcano Notes GEO.416 VOLCANOLOGY GEO.416 VOLCANOLOGY I. Physical Nature Of Magmas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4 Structural State of Silicate Melts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4 Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 Controls on Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5 Silica composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6 Volatiles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 Pressure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 Crystal content . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7 Bubble Content . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8 Yield Strength . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8 Specific Heat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8 Thermal Conductivity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8 Density . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 Electrical Conductivity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 Seismic Wave Velocities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9 II. Generation, Rise And Storage Of Magma . . . . . . . . . . . . . . . . . . . . . . . . . . 10 Nature of Crust and Upper Mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10 Heat Sources . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 Mechanisms of Melting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 Partial Melting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11 Segregation and Rise of Magmas Through The Mantle . . . . . . . . . . . . . . . . . . . . . . . 12 Rise of Magmas Through Brittle Lithosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13 Flow of Magma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 Flow Rates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14 Nature of Flow Regime . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 Flow Instabilities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15 High-Level Reservoirs and Subvolcanic Stocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16 III. Eruptive Mechanisms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18 Opening Of Vents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18 Mechanisms of Explosive Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18 Nature of the Gaseous Eruptive Column . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19 Bubble Nucleation And Growth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 Pressure Relations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20 Ejection Of Pyroclastic Material . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 Ejection Velocities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21 Eruption Energy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22 V. Lava Flows . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 Volume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23 Length and Thickness . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23Velocity of Flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 Discharge Rates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24 Physical Properties of Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 Temperature and Cooling of Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25 Morphology Of Lava Flows . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 Pahoehoe Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 External Structures of Pahoehoe Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26 Internal Structures of Pahoehoe Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27 Aa Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28 External Structures of Aa Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28 Block Lava . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 Internal Structures of Blocky Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29 Pillow Lava . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30 VI. Volcanic Domes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31 External Features of Volcanic Domes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31 Internal Structures of Volcanic Domes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32 VII. Products Of Volcanic Explosions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 Terminology and Classification . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 Origin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 Fragment Size . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33 Airfall Ash Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35 Dispersal . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 Structures Of Airfall Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36 Morphology of Ash Particles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 Pyroclastic Flow And Surge Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37 Relationship to Topography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38 Flow Units and Cooling Units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39 Components . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 Characteristics of Ash-Flow Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 Internal Layering . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40 Gas-Escape Structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 Textural Relationships . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41 Segregation of Crystals and Lithics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42 Temperature Effects . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42 Welding and Compaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42 Structures Related to Temperature and Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . 43 Classification and Nomenclature of Pyroclastic Flows . . . . . . . . . . . . . . . . . . . . . . . 43 VII. Laharic Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47 General Features . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47 Surface of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47 Basal Contact of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 Components of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 Grain-Size Distribution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 Grading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48 Fabric . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49 Origin of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49 VIII. Structures Built Around Volcanic Vents . . . . . . . . . . . . . . . . . . . . . . . . . 50 Cinder Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50 External Form . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50Internal Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 Maar Volcanoes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51 Littoral Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 Shield Volcanoes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 Icelandic Shields . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52 Hawaiian Shields . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 Galapagos Shields . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53 Composite Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 External Form . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 Internal structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54 Growth Sequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 Parasitic (adventive) Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55 IX. Craters, Calderas, and Grabens . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 Explosion Craters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 Collapse Craters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56 Classification of Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 Krakatoan Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57 Katmai Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 Valles Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 Hawaiian Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58 Galapagos Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 Masaya Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 Atitlán Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59 Cauldrons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 Volcano-Tectonic Depressions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 Resurgent Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60 X. Classification Of Volcanic Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 Nature of Vent . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 Styles of Eruptive Activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 Hawaiian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61 Strombolian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 Peléean Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 Plinian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62 Vulcanian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 Surtseyan Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63 Appendices . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 A. Pyroclastic Fall Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65 B. Pyroclastic Flow Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66 C. Pyroclastic Flow Deposit Characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67 D. Pyroclastic Surge Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 68 E. Pyroclastic Surge Deposit Characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69GEO.416 VOLCANOLOGY GEO.416 VOLCANOLOGY I. Physical Nature Of Magmas I. Physical Nature Of Magmas Magma is a completely or partially molten natural substance, which on cooling, solidifies as a crystalline or glassy igneous rock. It is usually rich in silica and capable of flowing under moderate differential stress. Magmas may carry rock fragments or crystals in suspension, and they normally contain gaseous (volatile) components in solution. Volcanic magmas fall within a strictly limited compositional range that reflects the physical and chemical processes responsible for their generation and differentiation. Our concern is the physical phenomena of volcanism, interpretation of which requires some knowledge of physical properties of magmas. Unfortunately, we have only a meager knowledge of liquid properties. Much of what is known can be explained in terms of the properties of Silicon (Si) and Oxygen (O) ions, which are usually the most abundant components. Si has a high charge (+4), small ionic radius (0.39 Å), and low coordination number with oxygen (4 oxygens surround each silicon, forming the corners of a tetrahedron). This results in strong ionic field strength and bonding with oxygen compared to other cations: Ca, Mg, Fe, Mn, Ti, Na or K. Al, which has similar but not as strong properties, plays a similar role to Si in both liquids and crystalline solids. Structural State of Silicate Melts Structural State of Silicate Melts Modern concepts of silicate liquid structure are based on the Zachariasen Model. The atoms are bonded by forces similar to those between atoms of crystals, but lack long range periodicity and symmetry. The magmas have silica (and alumina) tetrahedra linked (or polymerized) in three- dimensional networks in which (bridging) oxygen atoms are shared by two or more tetrahedra; the Si and Al cations are termed "framework cations." Other cations enter the melt in limited amounts as independent ions occupying positions between tetrahedra, and modify the basic structural framework and its physical properties; these cations, Ca, Mg, Fe, Mn, Ti, Na, and K, are termed "framework-modifying cations." The framework-modifying cations can be accommodated in amounts of up to about 20 cation percent before the basic framework breaks down into smaller geometric units. In breaking liquid continuity into smaller units, the framework changes from an extensive network of tetrahedra, all of which are linked by shared O atoms to smaller units with lower Si:O ratios until, when more than 66% of the cations are framework modifiers, the liquid consists of separate tetrahedra not directly linked to each other. Melt structure controls the physical properties of a magma. Viscosity is the most important of these properties, because it plays a role in factors controlling both the style of volcanic eruption and the physical nature of volcanic products. 5Viscosity Viscosity Viscosity is a fluid's internal resistance to flow. It represents the ratio of shear stress to rate of shear strain applied to a layer of thickness Z and permanently deformed in a direction x parallel to the stress. Mathematically, viscosity is expressed by: s = s o + h dm dt n , where s is the total shear stress applied parallel to the direction of deformation; s o is the yield strength of the fluid or the stress required to initiate flow; h is the viscosity, expressed in units called poises (dyne sec/cm 2 ); dm/dt is the gradient of velocity dx/dt or strain rate over a distance Z normal to the direction of shear; and, n is an exponent which has a value of 1.0 or less depending on the form of the velocity gradient. For many fluids, this expression describes a linear relation between the strain rate (dx/dt) and shear stress parallel to the direction of shear. If a shear stress greater than the yield strength (s > s o ) is applied, the resulting strain has two components: (1) elastic and recoverable; and, (2) viscous and non-recoverable. If a stress less than yield strength (s < s o ) is applied, the substance is deformed elastically and returns to its original form after the stress is removed. Some fluids do not require application of some initial force before they are permanently deformed by shear stress parallel to the direction of shear. Such fluids are said to exhibit Newtonian behavior when n equals 1.0 and s o equals zero. Highly polymerized or non-Newtonian fluids (known as Bingham liquids) have a finite yield strength that must be exceeded before they can be deformed permanently. In other words, Bingham fluids behave elastically until their yield strength is exceeded. Cooling and crystallizing magmatic liquids behave as newtonian fluids only until they contain approximately 20% crystals. Liquids with suspended solid particles may have a non-linear relation of shear stress to strain rate, for which the value of n is less than 1.0. Controls on Viscosity Various factors control magmatic liquid viscosity: composition (especially Si and volatiles), temperature, time and pressure, each of which effect the melt structure. Actually, the viscous behavior of complex silicate liquids, such as magmas, is difficult to predict, because no comprehensive theory explains the effects of major cations or temperatures of magmatic conditions. It is possible to estimate the viscosity of a magmatic liquid at temperatures well above liquidus temperatures (that is, temperatures at which only liquid is present) from chemical compositions and empirical extrapolation of experimental data on the linear relationship between h and temperature in simple chemical systems. The range of temperatures of naturally flowing magmas, however, is near or within the crystallization interval, where stress-strain relationships 6are not linear (that is, they are crystal-liquid mixtures and show Bingham behavior). Under such conditions, the only way to predict viscosities is by analogy with similar compositions investigated experimentally. Silica composition The strong dependence of viscosity of molten silicates on Si content can be illustrated by those of various Na-Si-O compounds: Na:Si:O h (poises) 0:1:2 10 10 1:1:2.5 28 2:1:3 1.5 4:1:4 0.2 The decrease in viscosity can be attributed to a reduction in the proportion of framework silica tetrahedral, and therefore, strong Si-O bonds in the magma. Temperature Temperatures of erupting magmas normally fall between 700° and 1200°C; lower values, observed in partly crystallized lavas, probably correspond to the limiting conditions under which magmas flow. Low temperatures characterize silica-rich rhyolite magmas, whereas the highest temperatures are observed in basalts. Magmas do not crystallize instantaneously, but over an interval of temperature. Few magmas, however, have a wide enough range of crystallization to remain mobile at temperatures far below those at which they begin to crystallize or much hotter than those temperatures. Temperature has a strong influence on viscosity: as temperature increases, viscosity decreases, an effect particularly evident in the behavior of lava flows. As lavas flow away from their source or vent, they lose heat by radiation and conduction, so that their viscosity steadily increases. For example: a) measured viscosity of a Mauna Loa flow increased 2-fold over a 12-mile- distance from vent; b) measured viscosity of a small flow from Mt. Etna increased 375-fold in a distance of about 1500 feet. The decrease in viscosity can be attributed to an increase in distance between cations and anions, and therefore, a decrease in Si-O bond strength. Time At temperatures below the beginning of crystallization, viscosity also increases with time. If magma is undisturbed at a constant temperature, its viscosity may continue to increase for many hours before it reaches a steady value. The viscosity increases with time results partly an increasing proportion of crystals (which raise the effective magma viscosity by their interference in melt flow), and partly from increasing ordering and polymerizing (linking) of the framework tetrahedra. 7Volatiles The solubility of gases in magmas varies with pressure, temperature and composition of both the gas and the magmatic liquid. Because the volume of a melt with dissolved gas is less than that of a melt and separate gas (vapor) phase, solubility increases as gas pressure increases. At constant gas pressure less than total pressure, any increased load pressure on the melt lowers solubility, because the volume of the melt with dissolved gas is greater than that of melt alone. Vapor pressure increases with temperature, so that solubility of any volatile component generally decreases with temperature, except possibly at high pressure. Consequently it is difficult to predict how volatile content of magma varies with depth. Nevertheless, it has been shown that at constant temperature, solubilities of water in magmas with different compositions are not significantly different. Nearly all magmas can contain more water or gases at depth than they can continue to hold in solution when they reach the surface. Basalts, however, normally contain less water than rhyolites simply because their temperatures are higher, and thus, as noted, lower gas solubility. Only limited data exists concerning the effect of volatiles (in particular F, Cl, S, H 2 S, SO 2 , CO, and CO 2 ) on magma viscosity. No doubt, the effect of dissolved water is to lower viscosity, the effect being greater for silica-rich than silica-poor magmas: Magma T (°C) h dry (poises) h wet (poises) Rhyolite (~70% SiO 2 ) 785 10 12 10 6 (5% H 2 O) Andesite (~58% SiO 2 ) 1000 10 4 10 3.5 (4% H 2 O) Basalt (~48% SiO 2 ) 1250 10 2 10 2 (4% H 2 O) Dissolved water disrupts the framework of linked Si and Al tetrahedra, but where such polymerization is already minor or absent, there is little effect. F and Cl are though to considerably reduce magma viscosities; in contrast, CO 2 increases polymerization, and therefore viscosity, in melts by forming CO 3 -2 complexes. Pressure The effect of pressure is relatively unknown, but viscosity appears to decrease with increasing pressure at least at temperatures above the liquidus. As pressure increases at constant temperature, the rate at which viscosity decreases is less in basaltic magma than that in andesitic magma. The viscosity decrease may be related to a change in the coordination number of Al from 4 to 6 in the melt, thereby reducing the number of framework-forming tetrahedra. Crystal content The effect of suspended crystals is to increase the effective or bulk viscosity of the magma. The effective viscosity can by estimated from the Einstein-Roscoe equation: h = h o (1 - RC) -2.5 8where h is the effective viscosity of a magmatic liquid, C is the volume fraction of suspended solids; h o is the viscosity of the magmatic liquid alone; and, R is a constant with a best-estimated value of 1.67. Bubble Content The effect of gas bubbles (vesicles) on the bulk viscosity of magmas can be variable, and depends on: (1) the degree of bubble formation (that is, vesiculation); (2) the size and distribution of bubbles; and, (3) the viscosity of the intervening melt. Exsolution of water increases viscosity, but the exsolved vapor is a very low viscosity fluid; in basaltic magmas, the bubbles may enhance the already low temperature and composition controlled viscosity. Rhyolitic magmas have high viscosities irrespective of the degree of vesiculation, and only effect of high bubble content will be to reduce mechanical strength of the melt. Yield Strength Yield Strength Most magmas have an appreciable yield strength, which shows a marked increase below their liquidus temperature. As yield strength increases, the stress required to initiate and sustain flow becomes greater, and the magma's apparent or effective viscosity is also increased. Specific Heat Specific Heat The specific heat (Cp) of magma, which is the heat required to change the temperature of the liquid 1 degree Celsius, is typically about 0.3 cal. gm -1 . The specific heat contrasts greatly with heat of fusion or crystallization, which is the heat that must be added to melt or removed to crystallize a unit mass that is already at a temperature where liquid and solid coexist. Heats of fusion are typically about 65-100 cal. gm -1 at 1 atmosphere. Consequently, about the same amount of heat is involved in crossing the crystallization interval, as in raising or lowering the temperature of the rock or liquid through 300°. Thermal Conductivity Thermal Conductivity Igneous rocks and liquids are poor conductors of heat. Thermal conductivity depends on two heat transfer mechanisms: (1) ordinary lattice or phonon conduction; and, (2) radiative or photon conduction. The former declines and the latter increases as temperature increases and the melt structure expands. For rocks, the two effects balance each other up to their melting range. At high temperatures, the thermal conductivity of mafic rocks normally declines at an increasing rate up to 1200 ° C, above which, radiative heat transfer increases as does total thermal conductivity. More silica-rich rocks show increasing thermal conductivity at lower temperatures. 9Density Density Magma densities range from about 2.2 gm cm -3 for rhyolite to 2.8 gm cm -3 for basalts, illustrating a close density-melt composition relationship, primarily reflecting the influence of higher concentrations of Fe, Mg and Ca cations in basalts. In contrast, magma density decreases with increasing temperature and gas content. These densities increase a few percent between liquid and crystalline states. The temperature dependence of magma density is given by the coefficient of thermal expansion, about 2-3 x 10-5 deg -1 for all compositions. The pressure dependence of magma density is given the compressibility or fractional volume change, D V/V, per unit of pressure. Compressibility increases sharply in the melting range from 1.3 x 10 -12 to about 7.0 x 10-12 cm 2 dyne -1 . Electrical Conductivity Electrical Conductivity Electrical conductivity, which is low in pure silica melts, increases with increasing abundance of metallic cations, especially alkali elements, and increases abruptly in the melting range. Seismic Wave Velocities Seismic Wave Velocities Compressional or P-wave velocities are about 6 km sec -1 up to the melting range, then decrease abruptly to 2.5 km sec -1 at higher temperatures. Shear or S-wave velocities are about 2-3 km sec -1 , which drop abruptly at melting temperatures. 10II. Generation, Rise And Storage Of Magma II. Generation, Rise And Storage Of Magma The subsurface processes by which magmas are generated and rise toward the surface are extremely complex. Before examining these processes, it is worthwhile to review what is known concerning the Earth's interior. Nature of Crust and Upper Mantle Nature of Crust and Upper Mantle Most of what is known concerning the Earth's interior comes from geophysical measurements, and concerns: (a) seismic wave velocities; (b) temperature; (c) density distributions; (d) heat flow; and, (e) mechanical properties. Seismic velocities increase with depth within the Earth, but show abrupt changes at several depths interpreted to represent discontinuities in the composition or structural state of minerals. The most notable discontinuities are: (a) Mohorovicic discontinuity (MOHO); (b) Low Velocity Zone (LVZ); and (c) Core-Mantle boundary The seismic velocities are closely related to the density r and the elastic properties (bulk modulus K and rigidity or shear modulus m ) by the following expressions: V p = { [K + (4/3)m ]r } V s = (m /r ) 1/2 The elastic properties are poorly known, but making certain assumptions, it appears that density increases to about 3.4 gm/cc at depths around 70 km, remains constant between 3.45 and 3.63 to the base of the Low Velocity Zone. Both pressure and temperature increase with depth. The temperature increase (6°/km) in the crust is consistent with an average heat flow of 1-2 x 10 -6 cal. cm -2 sec -1 , with the highest values associated with young crust. If temperature gradients measured in the crust are projected downward, they rapidly approach temperatures for beginning of melting in the mantle near the Low Velocity Zone. The transmission of shear or S seismic waves, however, suggests the absence of large amounts of liquid, so that the temperature gradients must diminish with depth. Heat Sources Heat Sources Existence of magma indicates that at some depth beneath the Earth's surface, temperatures must be high enough to induce melting. One major problem associated with understanding the generation of magmas is the source of heat necessary to cause melt production. It is believed that 11the major source of heat within the Earth is the radiogenic elements, principally K, U, and Th. These elements, however, are concentrated within the Earth's crust, and have extremely low abundances in probable mantle rocks, too low to yield through their radioactive decay the heat necessary to generate magmas. Moreover, it can be shown that the melting process scavenges these elements, and thus, depletes even more their abundances in the source region. Mechanisms of Melting A variety of models have been invoked to explain the source of heat required to induce melting within the Earth: (a) Stress Relief: Pressure on the source region is released during tensional or compressional deformation of the overlying rock column. (b)Thermal Rise to Cusp in the Melting Curve: Intersection of pressure- temperature conditions with the source rock melting curve under conditions where lowest temperatures on the solidus coincide with phase change boundaries. (c) Convective Rise : The source material rises by solid-state convection into a pressure-temperature regime appropriate for melting (d) Perturbation: A local decrease in thermal conductivity or density leads to heating or diapiric rise of the source material. (e) Mechanical Energy Conversion To Heat: Force required to move one rock surface over another without grinding and deformation converted to heat, because of thrust faulting, subduction, a propagating crack or flaw in the Earth's lithosphere, shear or Tidal energy dissipated in the solid earth. (f) Compositional Change: The addition or subtraction of material changes the rock composition to a new composition whose solidus lies at a temperature less than the ambient temperature. Partial Melting Rocks are a heterogeneous assemblage of minerals, and each mineral is characterized by a unique melting temperature. Melting commences at grain boundaries, usually where three crystals of minerals with the lowest melting temperatures meet. As melting progresses, channelways develop between grains. Temperatures probably never are high enough to completely melt the source rock, and only part of or some of the minerals melt. This process is therefore called partial melting. Because of mechanical constraints, it is generally believed that at least 1-5% melting is required for the melt to separate from the unmelted (refractory) solid (crystalline) material. Melting probably never exceeds 35% because of the gravitational instability of low density liquid with higher density refractory minerals. The composition of a partial melt (magma) depends on the melting conditions present in the Earth: (a) temperature; (b) pressure; 12(c) volatile content; (d) mineral composition of the source rock; and, (e) amount or degree of melting. Once gravitational instability sets in, the melt separates from the solid (denser) residuum. Depending upon where separation occurs, the magma may ascend through ductile (mantle) and/or brittle (crust) domains within the Earth. The manner in which magma rises differs between these two domains. Segregation and Rise of Magmas Through The Mantle Segregation and Rise of Magmas Through The Mantle Several mechanisms of magma rise through the mantle have been visualized. These processes include: (a) Deep Segregation: The melt forms along a dendritic network of joints and fractures in the zone of melting, and feeds into a smaller number of layer tributaries eventually forming a larger channel at higher levels. With melting concentrated along grain boundaries, melt migration is caused by a thermal or pressure gradient or by capillary effects. This migration the presence of a critical proportion of melt before solid/liquid separation occurs. Two factors which could provide the driving force following initial separation are: (i) pressure resulting from volumetric expansion on melting, and, (ii) the buoyancy of the liquid. Once the liquid has separated, it is unlikely that it maintains a temperature much higher than its surroundings, as it is cooled by adiabatic expansion and conduction to the wall rocks. If the liquid rises slowly through rocks that are below their melting temperature, the magma would crystallize quickly. Thus, magmas can only ascend once the temperature of their wall rocks have been elevated, and successive batches of magma must tend to follow paths of earlier bodies. (b) Diapiric Rise: A density reversal can lead to what is known as Rayleigh- Taylor instability in which lighter underlying material first collects in localized bulges under the heavier layer. The low density layer moves upward at an accelerated rate until it forms a steep sided plume or vertical density current. The rate of ascent , size, and spacing of plumes is a function of density differences, and the viscosity of the overlying rocks. Little or no separation of melt occurs in the zone of melting. Instead, the crystal-liquid mush rises and separation occurs at shallow levels. There again must be a delicate thermal balance between the diapir and its surroundings. Otherwise, it crystallizes. (c) Zone Melting: A body of magma rises by melting its roof, while it crystallizes on its floor. The zone of melting rises without actual movement of liquid and with little loss of heat. Heat used in melting is regenerated by release of latent heat of crystallization. It has been estimated that a body of magma 7 km thick starting at a depth could rise to within 8 km of the surface before crystallizing in about 1 million years. 13Rise of Magmas Through Brittle Lithosphere Rise of Magmas Through Brittle Lithosphere It is difficult to determine the level at which the lithosphere deforms by brittle fracture rather than by plastic flow - a depth represented by earthquake foci. There is strong evidence, in the form of individual and swarms of dikes, that large bodies of magma are tapped within the crust at a level where rocks can fail by dilational fracture. However, temperatures and pressures in the vicinity of large magma bodies are not normally consistent with purely brittle fracture. The manner in which magmas rise through the lithosphere may be: (a) Dilational Rise: This proposed mechanism by which magma may rise involves: (i) entrance of melt in fractures, and rise due to gravitational buoyancy; (ii) The fracture becomes extended vertically and/or horizontally along a plane normal to the minimum stress; and, (iii) The fracture closes behind the magma as it passes and pressure on the wall falls below the confining pressure, rebounding due to viscoelastic deformation. Such a mechanism may explain the limited duration of basaltic fissure eruptions and the apparent arrival of discrete batches. Many instances, however, exist where acid or volatile magmas have apparently risen as pipe-like intrusions with little or no evidence of horizontal deformation. The ability of a magma to rise through brittle lithosphere is usually explained in terms of depth and density contrast with the overlying rocks. If the pressure on the magma is equal to the lithostatic load of overlying rocks, the magma can rise to a level determined by the density contrast. At a depth of 50 km, the lithostatic pressure can exceed the pressure of a vertical magma column enough to segregate liquid and cause it to rise. If the heights to which magmas can rise is solely dependent on the depth to source and a density equilibrium, it would be expected that magmas with deep sources would erupt at higher elevations, and vice versa. This is obviously not the case as demonstrated by volcanoes of the Mexican volcanic belt. More important limitations to magma rise are probably the heat content, and rates of ascent and cooling, which in turn, depend on the size of the magma body. Another important factor is the stress regime, which governs the form of the intrusive bodies. The three basic magma stress regimes are: (a) least principal stress is horizontal (dikes); (b)least principal stress is vertical (sills); and, (c) the stresses (vertical and horizontal) are equal (pipes; random dikes and sills). At relatively high magmatic pressures or at shallow depths where vertical and horizontal stresses are low and about equal on the surrounding rocks, the magma conduits tend to be cylindrical. Thus, the form taken by a magma body may change drastically during its ascent. It is likely that near the surface, a cylindrical pipe is the most efficient form of conduit, because flow velocity increases and heat losses decrease as the horizontal section increases 14in size and becomes equidimensional. Thus, conduits tend to become centralized at the intersection of two or more fracture systems. (b) Non-Dilational Rise: As mentioned previously, there is ample evidence that some magmas have forcibly displaced rocks into which they have intruded, but others have made room for themselves by stoping or elevating the roof rocks. It is obvious that the critical elements are heat, and the manner in which the magma crystallizes, the shape and size of the body, and the volatile content of the magma. An excellent example of non-dilational rise is illustrated by the formation of diatremes, steep-sided, more or less cylindrical or funnel- shaped breccia pipes formed by penetration of crust by moderate- temperature, gas-rich magma (kimberlite and carbonatite). Two mechanisms may be capable of boring through the Earth's crust and creating diatremes: (i) Highly energized gases of deep-seated origin bore through the crust, opening channelways for the rapid ascent of magma; or, (ii)Explosive eruption is triggered by vaporization of heated groundwater propagated downward as pressure is released on progressively deeper gas-charged horizons. Flow of Magma Flow of Magma Knowing the rheological or fluid properties of magmas, we might be able to apply basic fluid dynamic principal to predict flow regimes of intrusive and extrusive magmas under various physical conditions. Unfortunately, a rigorous approach to our understanding of flow characteristics is not currently possible in the face of incomplete information about essential parameters of specific cases. Nevertheless, some insight into magma ascent processes may be gained by considering simple examples and approximations. Flow Rates The volumetric flow rate of a viscous fluid through a cylindrical channel under a constant pressure gradient is given by: Q = (RP r 4 )/8h L where Q is the volume flow rate in cm 3 sec -1 , P is the pressure drop in bars, r is the channel radius in cm, h is the viscosity of the fluid in poises, and L is the length of the channel in cm. Applying this relationship to a large (about 200 km 3 ) simple funnel-shaped magma chamber which is filled with basaltic magma (h = 300 poises) via a 3-km-long, 200-m-wide, cylindrical feeder pipe at its base and a pressure drop through the pipe of 1000 bars (1 kb/3.3 km), we find: Q = (3.14 x 1000 x 10 16 )/(8 x 300 x 3 x 10 5 ) = 4.36 x 10 10 cm 3 /sec or 3.76 km 3 /day. 15This simple calculation is important in that it illustrates that movement of large quantities of magma in short periods of time is entirely feasible. Nature of Flow Regime The type of flow imposed on a magma, that is, laminar or turbulent flow, is also of interest. For example, in the case of an initially heterogeneous magma, the liquid would become effectively homogenized by turbulence. The conditions that determine laminar or turbulent flow can be determined by calculating the dimensionless Reynolds number, Re, which in terms of average flow rate is given by: Re = (2r Q)/h r P where r is the density of the fluid. Turbulent flow occurs when Re > 2000. For the previous example, with r = 2.6 gm/cm 3 , Re = (2 x 2.6 x 4.36 x 10 10 )/(3.14 x 10 4 x 300) = 2.39 x 10 4 Hence, flow of the basaltic magma within the conduit would be turbulent. The higher viscosity of acid magmas, however, renders turbulent flow unlikely in these cases. Because the viscosity of magmas normally exceeds 10 3 poises and velocities are rarely greater than a few cm/sec, flow is probably laminar under most geologic conditions. It can be expected that the non-Newtonian characteristics of magma also have an effect on flow behavior. Because a certain yield strength must be exceeded before many magmas can be deformed by viscous flow, velocity gradients in the margins of a moving magma are likely different from those of more familiar liquids like water. Shear stress in the boundary of the moving liquid is greatest near a stationary surface and diminishes toward the interior. Thus, if viscosity is uniform throughout the entire flow width, then the velocity distribution is parabolic. But if heat is lost at the stationary boundary and the effective viscosity increases sharply with falling temperature, the flow profile is more arcuate. These different flow profiles reflect both the effect of falling temperature on both viscosity and yield strength of the magma. In many cases, it is likely that a zone of static liquid will form a layer between the moving liquid and its solid boundary. Heat transferred from a cooling magma to surrounding wall rocks also affects its behavior in other ways. Flow Instabilities When heat losses from the top or sides of a magmatic body cause a density difference in the liquid large enough to produce gravitational instability, the liquid overturns and free convection accelerates the rate of heat transfer. The onset of convection in an infinite horizontal layer of viscous fluid having an upper and lower surface is given by the dimensionless ratio of buoyant to viscous forces known as the Rayleigh number, Ra: Ra = (L 4 a T gb )/h K 16where L is the height of the layer in cm, a T is the coefficient of thermal expansion, g is the constant of gravitational acceleration (980 cm/sec), b is the vertical temperature gradient in K cm -1 , h is the kinematic viscosity (h/r ), K is the thermal conductivity of the magma in cal gm -1 K -1 , and r is the fluid density. Ra for a vertical tube heated from below is given by the same expression, except that L 4 is substituted by r 4 where r is the characteristic radius of the tube in cm. The critical Ra value above which convection begins is about 1700, approximately the same value calculated for magmatic bodies of most common shapes. For a magma body of given size and viscosity, the principal variable is thermal gradient, b , a function of heat loss to the top or sides of the magma body. For Ra < 1000, transfer of heat is predominantly by conduction; steady convective heat transfer sets in at approximately Ra > 10000, and strong eddying motion is attained when Ra = 100000. Bodies with thickness or radius greater than 10 m are likely to convect if their heat losses are those that would be expected at shallow crustal depths (10 -5 to 10 -3 cal cm -2 sec -1 ). Clearly, the larger the magma and the lower its viscosity, the more likely convection occurs, but quite small bodies having high heat flux values, should also be quite unstable. High-Level Reservoirs and Subvolcanic Stocks High-Level Reservoirs and Subvolcanic Stocks The erosion of extinct volcanoes reveals the presence of simple and multiple stocks of medium- to coarse-grained rocks. Generally, the stocks are 1- to 10-km-wide, circular to oval in cross-section, and grade upward into a maze of inward dipping sills, steep radial dikes, and cone sheets. Most of these intrusive rocks have made room for themselves by stoping rather than forcible intrusion. There is good evidence that these intrusive bodies were volcanic reservoirs, because compositional features of erupted materials indicate that most magmas tended to reside and equilibrate in such shallow reservoirs prior to eruption. Other than what we see within deeply eroded volcanoes, however, little is known concerning the volcanic reservoirs beneath active volcanoes, except what is indicated by geophysics: (1)Seismic methods: These methods have been used to detect large magma bodies at depth because of the inability of the Shear or S seismic waves to be transmitted through liquids. The distribution of earthquakes generated within or directly below a volcanic structure may delineate: (a) the boundaries of intrusive bodies, and (b)the possible movement of magma within the subvolcanic plumbing system. For example, a three-dimensional distribution of earthquake foci surrounds an aseismic zone, which may represent one or more bodies of magma beneath Kilauea. Several types of earthquakes of volcanic origin are recognized according to the location of their foci and the nature of earthquake motion: (a) A-type volcanic earthquakes: These earthquakes take place in and beneath volcanoes at places deeper than 1 km, generally in the range from 1 km to 20 km. They are generally less than 6 in magnitude. (b) B-type volcanic earthquakes: These earthquakes originate usually in and adjacent to active craters at extremely shallow depths. The 17magnitudes are generally extremely small. The earthquake motions consist mainly of vibrations with periods in the range of 0.2 sec. to 1.0 sec. (c) Explosion earthquakes: The maximum amplitude or magnitude of the earthquake has a close relationship with the intensity of explosive eruption and is approximately proportional to the kinetic energy of the eruption. The earthquake motions show a predominance of longer wave length as compared with those of the A-type volcanic and tectonic quakes. The associated detonations or air vibrations of explosive eruptions are remarkably strong. (d) Volcanic tremors: Earthquakes take place incessantly or continuously with a short interval, such as every several seconds, so that motions are recorded continuously. These earthquakes may originate from extremely shallow positions in or near the crater, or at deep levels (20-30 km at Kilauea). Various wave forms are found in volcanic tremors, including surface waves of Rayleigh and Love type. (2) Gravity Measurements: Precise gravity measurements may also reveal the presence of an anomalous mass of magma at depth, and provide a means of constructing subsurface structural models. Gravity surveys have shown that the Hawaiian volcanoes have crudely cylindrical cores composed of dense rock only a few km below their summits. Gravity measurements have also suggested the presence of large batholith-size, low-density bodies of magma or intrusive rock beneath many large calderas. They also indicate that Cascade volcanoes lie within grabens, or down-dropped tectonic blocks, underlain by similar subvolcanic intrusions. (3) Infrared Radiometry: This technique is used to detect the presence of bodies of rock or magma at elevated temperatures. (4) Tiltmeter Measurements: Precise leveling and tilt measurements have been used to detect deformation caused by the intrusion of magma into shallow levels. Such measurements have been used to estimate the depth and geometry of the intrusions, because they provide precise information concerning the horizontal as well as the vertical components of movement. 18III. Eruptive Mechanisms III. Eruptive Mechanisms Opening Of Vents Opening Of Vents Rare observations indicate that during the initial phases of a volcanic eruption: (i) the fractures through which magma reaches the surface represent planes of dilation propagated ahead of slowly rising magma; (ii) the appearance of lava is preceded by a mild release of steam or heated groundwater; and, (iii) eruption typically involves extrusion of magma that is relatively rich in gas. The strength, porosity and water content of near-surface rocks, shape and dimensions of the vent, and the physical properties of magma have a greater influence on the eruptive behavior than the depth of magma origin. Few explosive events are singular in nature, but rather represent an erratic succession of surges. Magma does not reach the surface unless it is sufficiently heated to remain fluid and to penetrate the overlying barrier of cold rocks and groundwater. In order for these conditions to be met, it appears that a minimum conduit width and flow rate of magma within the feeder dikes is required. The final ascent of magma to the surface is neither sudden nor violent, but rather is a steady process that accelerates after the surface. The accelerated discharge may be due to: (a) reduced resistance to flow; (b) reduced density caused by expansion and vesiculation; (c) educed heat loss to surrounding rocks; and, (d) increased temperature resulting from shear heating adjacent to dike walls. The spacing and duration of eruptions seems controlled by the rates of stress accumulation in the lithosphere. Eruptions cease not because of a lack of magma, but due to a reduction in pressure. Mechanisms of Explosive Eruptions Mechanisms of Explosive Eruptions All explosive eruptions involve the sudden release of energy by gas under pressure, but the way gas expansion acts on magmas varies widely. The explosivity of a volcanic eruption does not correlate directly with either volatile or silica content of the magma alone: the lowest is in those of olivine basalts, but highest in those of basanites and lamprophyres. The major factors which determine the explosivity are: (a) the rate of gas expansion, and, (b) the manner in which expansion occurs. These factors, in turn, depend upon the viscosity of the magma, and the way in which they vesiculate. The degree of vesiculation and gas expansion may vary throughout an eruption. Following a period or repose, initial eruptions usually therefore involve a gas-rich magma. Thereafter, the volatile content declines as gases escape to the atmosphere, and viscosity increases as more gas-poor magma is tapped. Low-density gas, either juvenile (magmatic) or meteoric (groundwater), concentrates in the upper parts of the plumbing system or reservoir by diffusing through a narrow boundary layer, through convective processes or by vesiculation and rise of 19bubbles. Once a magma becomes saturated, it may rise and reach a level at which the pressures exerted by the overlying rocks are low enough to permit vesiculation. Expansion accelerates the rise of magma, so that the pressure of the overlying rock column is reduced at a faster rate, and eruption ensues. This process by which a reduction of lithostatic pressure allows an increase in exsolution of gas from the magma is known as "second boiling". Vesiculation could also be initiated by convective overturning of an density-stratified magma, or by injection of hotter magma (remember that, in both cases, a resulting temperature increase decreases gas solubility). In most cases, the initial phases of eruption result in the ejection of gases and disrupted magma or ejecta with in a gas-charged cloud or eruption column. Nature of the Gaseous Eruptive Column Nature of the Gaseous Eruptive Column To understand fully eruption mechanisms it is useful to examine the characteristics of the eruption column and how it varies as magma reaches to the vent: (a) Temperature Relations: Exsolution and expansion of gas significantly cools magma as it rises. If there is good thermal equilibration between the magma and gas, the extent of cooling can be very great, e.g. there can be 300°C cooling of a vesiculating basaltic magma, if it expands adiabatically from the pressure at which gas exsolution begins. The temperature of the gas is largely dependent on the proportion of the two phases, and the efficiency of the heat exchange. The latter is strongly dependent on size because only ejecta or magma fragments less than 5 mm can attain thermal equilibrium with the gas during an eruption; silicate particles therefore account for most of the heat. If the source of the gas is meteoric water, the heat used to flash the water to steam tends to buffer the temperature eruption at around 100°C. As the eruption column emerges from the vent, it continues to cool as it expands and mixes with air. (b) Density Relations: The density of the eruptive column influences its capacity to carry fragments suspended in the gas stream. The smaller particles are subject to drag forces larger than their inertial forces, and thus, have lower terminal velocities so that they behave like gas particles. Particles less than 0.1 mm in diameter have so low terminal velocities compared to the velocity of the gas stream, that they contribute to the effective density and viscosity of the eruption column. A greater proportion of fine particles therefore enhances the ability of the eruption column to support large clasts or fragments. (c) Viscosity Relations: A marked increase in magma viscosity occurs as a result of falling temperature and reduced water content during eruption. As a consequence, there is a slower expansion rate of bubbles as the magma approaches the surface. Conversely, the increased proportion of gas lowers the overall viscosity if the gas phase becomes large enough to be continuous. 20Bubble Nucleation And Growth Bubble Nucleation And Growth In order to understanding the mechanisms of explosive eruptions, it is useful to consider the manner in which gas exsolves from the magmatic liquid. Even in the most viscous magmas, the rate of bubble nucleation is very high. In order to evolve and grow, gas bubbles must reach an initial size that balances the surface tension (s ) of the magma at the gas-liquid interface. The pressure of gas inside the bubble acts over a cross-section P r 2 , and is balanced by surface tension around the circumference of its walls in the same cross-section (2P rs ). Therefore, the gas pressure must exceed a value of 2s /r before it can expand. Stable micron-sized bubbles can form if the gas pressure is greater than 6 bars (dry) or less (water-saturated). Phenocrysts (large suspended crystals) accelerate vesiculation because bubbles that nucleate on the crystal surface require less volume to reach a given radius. The surface tension at a gas- liquid interface increases with falling temperature, but may be offset by dissolved water. The exsolution of water vapor increases surface tension to different degrees in different magmas, which may explain why bubbles tend to expand intact in some magmas but coalesce in others. Exsolution and expansion of dissolved gases ultimately leads to disruption of the coherent magmatic liquid. Pressure Relations The principal factor controlling the violence of explosive eruptions is the magnitude of residual gaseous phase, when the magma approaches the surface. There are four components of pressure in the vesiculating magma: (a) the pressure of the overlying magma column: (r gh) (b) the pressure required to drive the magma through the conduit: P = 12Vh h/r 3 in cylindrical conduits P = 12Vh h/w in fissure conduits where V is the magma flow velocity, h is the length of the conduit, and r is the conduit radius or w is the fissure width. (c) the pressure required to overcome surface tension: The essential condition is the relationship between gas pressure in bubbles to the strength of the surrounding liquid. The strength of a vesiculating magma may be determined by the bubble density: when the proportion is low, it is an important factor, but as the proportion increases, surface tension becomes important. The force of surface tension acting around the circumference of each bubble exerts a pressure over the cross-sectional area of the bubble, so that the total pressure from surface tension through the vesiculated liquid is: P = 2n 2/3 s /r 21where n is the number of bubbles per unit volume and r is their average radius. The excess pressure of the gas phase, D P, exerts a force per unit of cross-sectional area of vesiculating magma, and must be greater than: D P > 2n 2/3 s /r + t where t is the critical tensile stress of the magma. For porosities greater than 50 percent, this excess pressure need only be a bars in order for fragmentation of the magma to occur. (d) The pressure required for the bubble to expand against the viscous resistance of the surrounding liquid: P = 4h / r (dr/dt), where (dr/dt) is the expansion rate of the bubbles. This pressure, which varies between 10 -2 bars and several hundred bars, is strongly dependent on magma viscosity. In a fluid basaltic magma, a bubble with a 1 cm radius can grow radially at a rate of 0.5 mm/sec, more than enough to accommodate gas expansion at low pressure, but in viscous magmas, the expansion rate is two to three orders of magnitude slower and the pressure buildup is greater. The final sizes and gas pressures of bubbles are mainly a function of magma viscosity: the effect of increased viscosity during exsolution arrests expansion when the volumetric ratio of gas to liquid is between 3:1 and 5:1. The first and second pressure components decrease as the magma rises and expands, whereas latter components are small. After the magma has vesiculated to the point that it behaves as a compressible fluid, i.e. the gas forms a continuous phase in which silicate liquid is carried in suspension, the second component, the dynamic pressure, becomes dominant. Ejection Of Pyroclastic Material Ejection Of Pyroclastic Material As mentioned previously, the ability of the eruption column to carry in suspension and eject fragments of disrupted magma is determined by the column density. The nature of ejecta and the manner in which it is thrown out of the vent during eruption depends on their origin: (a) primary material derived from the magma, or (b) lithic fragments derived from conduit walls, with most plucked from the sides of the vent but some brought from deeper levels. The principal difference in behavior of these fragments is that the primary magmatic fragments are part of the moving gas stream, whereas the accidental blocks are accelerated from rest. Ejection Velocities The muzzle velocities of ejecta depend on the size and settling velocity of fragments in the gas stream. The ejection velocity is the difference between the velocity of the gas stream and the velocity with which fragments would settle under static conditions. The minimum ejection velocity 22can be estimated from the maximum distance that blocks of a given size travelled from the vent to impact: R = V 2 sin2J/g where R is the distance, V is the initial velocity and J is the ejection angle. R has a maximum value when J = 45°. The ejection angle is seldom as low as 45°; ejection angles tend to be 80° or more above the horizontal, and increase with depth to the focus of explosions. Velocities calculated with this expression are less than the actual ejection velocity, because as soon as a block leaves the effect of the gas stream, air resistance reduces its range, especially when it is small and has little momentum. For a given velocity, moreover, the ejection distance varies directly with the mass of the block, and inversely with its drag coefficient and cross-sectional area. The drag coefficient varies with the shape, surface roughness, and velocity of block, and with the viscosity and density of the atmosphere. For a given initial velocity, large blocks travel farther than smaller ones, because their inertia is higher, and momentum is less retarded by air resistance. Below a few centimeters diameter, fragment movement is strongly retarded by wind and thermal currents. Estimated ejection velocities are on the order of 500-600 m/sec. Lower velocities are produced by the convective rise of warm air and gas. These currents, which are capable of carrying only fine dust, may reach great heights above the volcano, but velocities rarely exceed a few 10's of m/sec. The heights to which the eruption cloud rises therefore is related to: (a) the vent radius; (b) the gas velocity; (c) the gas content of the eruption; and, (d) the efficiency with which thermal energy is converted to potential and kinetic energy during interaction with the atmosphere. In general, large eruption clouds that reach high attitudes are produced by large eruptions of fine particulate material. Eruption Energy The energy release during a volcanic eruption is a summation of varied and often offsetting forces: (a) heat energy contained in the solid and fluid products; (b) heat and mechanical energy required to heat subsurface rocks and vaporize meteoric water; (c) mechanical energy expended by magma and gas expansion; and, (d) work done against gravity during ascent of the magma. 23IV. Lava Flows IV. Lava Flows Lava flows are the products of extrusion of a coherent magma body onto the Earth's surface. The external forms and internal structures of lava flows are the result of both the physical properties of the magma and the external environment in which extrusion takes place. The principal physical property that determines the nature of the lava flow is the magma viscosity, which is itself influenced by both the chemical composition of the magma and its temperature. The rate of magma supply to the flow is also important. The external environment includes the steepness of the slope on which the lava is deposited, and the presence or absence of water and/or ice. Volume Volume Basalts are not only the most abundant lavas, but they are also the most voluminous. Ultrabasic lavas are rare, and the abundance of andesitic, dacitic and rhyolitic lavas decreases as the magma viscosity increases with increasing silica and alkali content. The volumes of most historical lava flows are generally measured in the 10ths or 100ths of cubic kilometers. Some of the largest known lava flows include: (a) 1669 Mt. Etna lava - » 1 km 3 , (b) prehistoric McCarty Flow (New Mexico) - » 7 km 3 , and, (c) 1783 Laki basalt flow (Iceland) - » 12.2 km 3 . All of these lavas are basaltic; siliceous rarely exceed 1 km 3 , with individuals some times only a few square feet in area and a few inches thick being known. Length and Thickness Length and Thickness Because siliceous magmas are usually more viscous than basic ones, siliceous lavas tend to be the shortest and thickest of all flows. Some lava flows are formed by a single gush of liquid spreading as a single unit. More frequently, it is found that repeated gushes of liquid have given rise to intertonguing layers known as flow units. Subaqueous flow tend to remain fluid longer than terrestrial flows because with increasing depth of water, exsolution of volatiles is suppressed and viscosity remains high due to dissolved water. (a) Basaltic Lavas: Fluid basaltic lava flows in Hawaii extend for more than 35 km with an average thickness of 5 meters. Some Icelandic basalts can be traced 80 km, whereas several Columbia River plateau basalts extended for mor than 100 km from their source vents. One Columbia River basalt flow has been traced over an area of 130 X 240 km, and has a thickness between 30 and 50 meters. The length of lava flows is determined largely by the magma effusion rate. A high effusion rate, where lava spreads rapidly from the vent, usually results in a single flow unit. A low effusion rate, in contrast, results in lavas of limited extent that pile up layer on layer. It appears that basaltic lava flows 24that originate from fissures spread for distances that are roughly proportional to the third power of their thicknesses. (b) Andesitic Lavas: Andesitic flows generally have thicknesses of up to 30 meters, and are usually 5 to 15 km in length. Pyroxene andesite lavas typically are more extensive than those of hornblende andesite. Hornblende andesite lavas tend to form short stubby flow that have the form of domes. (c) Dacitic and Rhyolitic Lavas: Siliceous lavas are short and thick. Few of these lavas travel more than 1 or 2 km, and many come to rest on 30-40° slopes. Some of the largest known siliceous lavas include: (a) The Big Obsidian flow (Medicine Lake) - » 1 km long, (b) Glass Mountain Obsidian Flow - > 5 km in length, and, (c) Ring Creek dacite - 27 km long and up to 250 km thick. Velocity of Flow Velocity of Flow The flow velocity of lava flows depends on a number of different factors: (i) rate of effusion, (ii) magma viscosity, (iii) volume of magma extruded, (iv) magma density, and, (v) the slope and nature of the channel in which it flows. As expected, flow velocity diminishes with distance from the source. A pronounced velocity gradient exists within lava flows, extending from the middle toward the top, bottom and sides. Without a surface crust, the fastest movement occurs in the upper and middle parts of the flow, but once a crust forms, the fastest-moving part moves increasing downward into the lava. Some typical flow velocities are: (a) Basaltic Lavas 30-60 km/hr Hawaii 8-75 km/hr Vesuvius (b) Siliceous Lavas usually on the order of 10's or 100's of meters/hour. Discharge Rates Discharge Rates The discharge rate of lava flows from the volcanic vent depends principally upon the fluidity of the magma, and size and dimensions of the conduit. Like flow velocity, discharge rates decrease during the course of an eruption. Some typical basalt discharge rates are: (a) 1947 Hekla - 75000 to 1250 m 3 /sec, (b) 1887 Mauna Loa - 5 million m 3 /hr, and, (c) 1946 Parícutin - 2 to 6 m 3 /sec. The discharge rates of intermediate and siliceous lavas are generally much lower than those of basalt, but there are notable exceptions: (a) Sakurajima - 1666 m 3 /sec, and, (b) Santorini - 45000 m 3 /sec. 25Physical Properties of Lavas Physical Properties of Lavas Temperature and Cooling of Lavas Most lavas are erupted at temperatures below their beginning of crystallization, and only rhyolitic obsidians are aphyric, or free of crystals. Because of their low thermal conductivity and high specific heat, most lavas are well insulated and cool slowly. Relatively little cooling takes place through most of the course of the flow, especially if the eruption temperature is greater than 1100°C. The principal heat loss of a lava is through radiation from its surface. This can be expressed by the Stefan-Boltzmann equation: Q = sT 4 , where Q is the energy radiated per cm 2 /sec, T is degrees Kelvin, and s is the Boltzmann constant (5.67 X 10 -5 ergs/sec/cm 2 /deg 4 ). Because of the 4th power temperature relation, a small amount of cooling greatly reduces the radiative heat loss. Only a minimal amount of heat may be conducted to the air or ground, as indicated by: Q = 2K(T s -T o )(t/(Pa )) 0.5 , where Q is the heat flux per unit time t, K is the conductivity of the ground, a is the thermal conductivity, T s is the surface temperature, and T o is the initial temperature of the ground. Owing to the low thermal conductivity and thermal diffusivity of soils and rocks, heat losses due to conduction are only a degree or two per hour. Lava temperatures can be measured with (i) an optical pyrometer in which the color of incandescent lava is compared to that of a glowing filament; (iii) a sheathed thermocouple, or (iii) infrared techniques. A rough estimate of lava temperature (°C) may also be obtained from the color of the flowing magma: brownish-red 500-650° dark red 650-800° bright red 800-1000° orange 1000-1150° yellowish-white 1150-1300° Viscosity There are very few measurements of the viscosity of flowing lavas, but this property may be estimated from the relation: h = gr sinAd 2 /3V 26where V is the mean velocity, g is the acceleration of gravity, A is the slope angle, d is the depth of the flow, and r is the magma density. The denominator, 3V, is appropriate for a broad sheet, whereas 4V is typically used to model narrow channels. The viscosities estimated from this relation are low, because the velocities measured at the surface are greater than the mean velocity of the flow. It is also possible to estimate lava viscosities from surface wavelengths of ripples in the lava crust using: h = 2.61rl 1.5 With falling temperature and increasing crystallization, lavas become increasingly non-Newtonian, and therefore require greater shear stress before flowing. This change in the viscous behavior of the lava accounts for flow fronts and levees ceasing to flow laterally even though slope angles may be great enough. Morphology Of Lava Flows Morphology Of Lava Flows Lava flows exhibit a variety of morphologies that depend on the magma viscosity and the external environment. Several types of lava flows are recognized to occur in lavas of different bulk composition. Pahoehoe Lavas These lavas are characterized by smooth, billowy, ropy or entrail-like crusts of quenched lava. Based on external form, various subtypes of pahoehoe lava can be distinguished: (a) Massive: The lava crust is about 3 to 15 m thick, and smooth over large areas. (b) Scaly: The lava consists of many small lobes or flow units that overlap like fish- scales. These units, sometimes called pahoehoe toes, may be 3-30 m in width and up to 30 cm thick. (c) Shelly: This very frothy lava has a minutely spinose sharkskin-like surface. Locally, a ropy or corded surface develops when the fluid magma moves beneath a thin, partly congealed crust, causing it to wrinkle and fold either convex downstream or in parallelism with the flow direction. (d) Slabby: These pahoehoe lavas are characterized by broken crusts, forming slabs a few meters across and a few cm thick. External Structures of Pahoehoe Lavas These lavas types, depending on magma viscosity, may show a variety of small-scale surficial structures that include: (a) Lava Coil: These structures, which typify Shelly subtypes, consist of coiled, rope-like strips of magma crust, a few cm to about a m in diameter and 5-30 cm in height. The coils develop along shear zones between relatively stationary and adjacent blocks, being moved by undercurrents. (b) Lava Blister: A mound of continuous lava crust, a few mm to a m in height and width, caused by the accumulation of gas beneath the lava surface. (c) Tumulus: This dome-shaped structure, resembling an ancient burial mound and typically having an oval ground plan, forms as a result of upwelling of 27magma beneath a fairly thick lava crust. A tumulus forms when the lava below the crust is obstructed downstream. The tumulus may be up to 50 m in length, and is locally 6- to 10-m-high. The tumulus surface is similar to that of the surrounding flow, except that it is generally cracked radially. Lava often rises in the cracks to form either small unrooted lava flows, or bulbous mounds, up to a few m in height and width, called Squeeze-Ups. (d) Pressure Ridges: These transverse, convex downstream, ridges represent lava crust which has been heaved up into elongate mounds as much as 0.8 km long and 50 m high. These ridges form as a result of the flow crust being pushed against some obstacle by continued movement from behind. Elevation of the crust into an anticline is aided by the hydrostatic pressure of liquid beneath the crust. In some cases, continued movement of the lava results in the overturning of the fold with the gentle slope on the lava source side and a steep slope away from it. Locally, the folded crust breaks and slides forward over the steep side of the ridge, forming a thrust fault. The crust of pressure ridges is generally broken, and many of the ridges consist of a heap of variably oriented blocks. (e) Hornitos, driblet- and spatter cones: These are small mounds or chimney-like spires that are built over eruptive vents or more commonly cracks in the lava crust (rootless vents). They are formed by the discharge (locally explosively) of clots of lava that adhere to earlier clots to form a pile of welded ragged- surface fragments in a deposit called agglutinate. (f) Lava tree molds and casts: These are molds formed when lava flows around standing trees. After flow level subsides, a hollow cylindrical column is left by the carbonation of plant material. Internal Structures of Pahoehoe Lavas In addition to these external features, pahoehoe lavas may exhibit a number of internal structures which include: (a) Flow units: formed by the intertonguing of lava streams derived from the same flow. (b) Columnar Jointing: Contraction, the result of thermal stresses within the cooling lava, produces fractures that are propagated in a plane normal to the direction of cooling. These fractures bound 5- or 6-sided, polygonal columns that develop perpendicular to the cooling surface. The columns, which vary from 5 cm to >3 m in width, are typically straight and have parallel sides, but some may be curved. Throughout individual flow units, columns may variable considerably in dimension and cross-section, but a three-fold subdivision is typically recognized: (a) upper colonnade (b) entablature (c) lower colonnade. Invariably, the columns are cut by cross-joints, some curved upward or downward as ball and socket joints. Discontinuous cooling leads to the development on the sides of the columns of chisel marks which mark the position of isotherms during cooling. Although columnar joints are common in all types of lava flows, it should be noted that they also characterize some 28pyroclastic ash-flows that have been emplaced at temperatures high enough for fragments to become welded together. They are also conspicuous in some subvolcanic dikes (where they occur normal to the dike walls, stacked like firewood when exposed by erosion) and in intrusive necks like Devil's Tower. (c) Lava Tubes: These structures, which range from a few cm to 30 m or more in diameter and from a few km to 20 km in length, develop by the flow of molten magma within a confined interior channelway. In the upper parts of a lava flow, migration of magma eventually becomes restricted to these channelways. Fast-moving flows are characterized by relatively straight lava tubes, whereas slow-moving flows tend to contain meandering and branching tubes. If lava drains out of the tube before complete solidification, it leaves strandlines on the tube walls. In cross-section, lava tube walls are marked by concentric layers of congealed lava. Completely filled tubes show concentric bands of vesicles, platy joints parallel to the walls, and/or radiating joint columns. Lava stalactites and stalagmites may form by dripping of still-fluid lava from the tube ceiling; some may consist of sulfate minerals or opal. (d) Pipe vesicles and spiracles: These gas cavities are formed when lava passes across wet ground, generating steam. The steam bubbles rise into the lava and form lines of vesicles or small tubes, usually less than 0.5 inches in diameter. If the upper end of the gas tubes are bent in the direction of lava movement, they are called pipe vesicles, and have been cited as possible indicators of flow direction. Where the steam bursts upward into the lava, it explosively creates an irregular, up to 10 m diameter, cylindrical opening called a spiracle. The spiracle generally terminates within the flow rather than extending through it, and may contain mud blown up from the underlying ground. Aa Lavas Aa lavas are characterized by surfaces that are a jumble of rough, clinkery and spinose, fragments, small chips to blocks measuring meters, and grade downward into massive lava. Based on external form, various subtypes of aa lava can be distinguished: (a) Aa Rubble flow: The lava crust consists of small, loose and semi-detached fragments. (b)Aa Clinker flow: The lava crust consists of loose and semi-detached fragments that measure more than several cm in diameter. (c) Furrowed aa flow: The lava is intermediate between aa and pahoehoe, with a very rough ropy surface that is locally arborescent. Aa lavas flow like a caterpillar tread, dumping talus over the snout and then overriding their own debris. Hence, they consist of a central massive part between fragmental top and bottom. External Structures of Aa Lavas Aa lavas types, depending on magma viscosity, may show a variety of large-scale surficial structures that include: 29(a) Lava Gutters: These channels develop when faster-moving parts of the flow drain away from slower-moving parts and flow bottom as the supply of lava diminishes or stops. (b) Lava Levees: These longitudinal ridges develop by accretion of lava on the slower-moving parts or flanks of the flow, generally bounding the central gutter. (c) Lava Lobes: These features represent lava tongues that have generally developed along flow margins after the levee is breached. (d) Accretionary Lava Balls: These structures form, like snowballs, by the rolling up of solid fragments, either clinker or chunks, derived from the walls of the flow channels, and typically range in diameter from a few cm to 3 m. Block Lava These lavas, which have surfaces covered by angular fragments, differ from aa in that the fragments have more regular forms and smoother faces. The surface blocks often approach cubes in form. Blocky lava flows form from more viscous lavas than aa flows, with the angular blocks formed by breaking up of the partly to wholly congealed upper part of the flow as still-mobile magma moves beneath the thick crust. These flows are typically thicker that aa lavas (8-35 m thick), and fragmental material, which may constitute the entire thickness, makes up a greater proportion of the flow than aa. The surfaces of blocky lavas are generally very irregular, with many hummocks and hollows, often 3-5 m deep. Internal Structures of Blocky Lavas Blocky lavas display a number of characteristic internal structures which, in addition to composition, may allow them to be distinguished from aa lavas: (a) Ramp Structures: The high magma viscosity results in a large amount of internal shearing. Movement along the ground is retarded by friction, whereas moving liquid higher up in the flow tends to separate into a series of sheets that slip over each other like a deck of cards. Movement of the lava sheets is predominantly parallel to the underlying surface. When solidified, these sheets may be very thin (few cm), and are defined by platy joints. Near the flow front, the extremely high viscosity of the magma may cause the shear planes to bend sharply upward. Ramps may be formed when local movement upthrusts portions of the flow along the shear planes (b) Lamination: These structures are formed by the upward bending of flow planes and shear planes, often distinguished by different degrees of crystallinity. These laminations may form antiforms and synforms, the limbs of which may become crumpled. (c) Spines: These structures form when a massive central part of the flow is projected up into the fragmental portion of the flow, or even extended about it. (d) Auto-Breccia: These deposits represent brecciation of the flow resulting from shattering of the very viscous lava due to stress related to flowage. 30Pillow Lava These flows are subaqueously extruded lava marked by bulbous forms. Pillow lavas may form by the discharge of lavas into rivers, lakes, ponds or under glaciers, as well into oceans. The pillow structures result from the protrusion of elongate lava lobes, which detach from and fall down the moving flow front. Lava pillows are often confused with pahoehoe toes, but the former have several distinguishing characteristics: (a) Pillows are rimmed by chilled glass selvedges, formed by rapid cooling of lava by the surrounding water. (b) Pillows vary from a few cm to several m in diameter, and are generally spheroidal, ellipsoidal, or may be flattened in cross-section. (c) Pillow tops are usually convex upward, whereas their undersurfaces may be flat, concave upward or project downward between the underlying pillows. (d) Where gas cavities are present, these structures tend to be located within the upper part of the pillows. Pillow lavas are locally found in association with several other types of subaqueous volcanic deposits: (a) Hyaloclastites (Aquagene Tuffs): These deposits are made up of broken pieces of glass, formed by brecciation as a result of drastic chilling of the fluid lava. (b)Pillow Breccia (Aquagene Breccia): These deposits are similar to hyaloclastites, but are dominantly composed of pillows and pillow fragments. 31V. Volcanic Domes V. Volcanic Domes Over or near the vent, extremely viscous lava tends to pile up into steep-sided heaps of molten rock, known as domes or tholoids. Some domes also result from the bodily upheaval of material filling the upper part of the conduit. This semi-solid to solid material is pushed up like a cork from the neck of a bottle, and are referred to as plug domes or belonites. Where the heap originates from outpourings of viscous lava, it may grow by addition of lava either internally or externally. Those domes that form by addition of lava through some form of extrusion through an opening in their crust, generally at the crest, are called exogenous. More commonly, lava squeezed up through the vent distends the mass above, so that these domes are called endogenous. Most domes are composed of rhyolite, dacite or trachyte magma. Andesite domes are less common, and basaltic domes are extremely rare. The size of domes varies greatly. Some are only a few meters across and a meter high, whereas the Mount Lassen dome is over 1 km across at its base and over 600 m high. Most domes are broader than high. In plan, domes are more or less circular, or very short ovals. Rarely are they elongate except as a result of extrusion of lava through a fissure vent. Where extrusion occurs as concentrations along a linear fissure, a row of clearly separate and independent domes, sometimes with overlapping bases, may grow simultaneously. More commonly, domes grow successively along a fissure. Most domes are short-lived features, because they are commonly destroyed by collapse partly due to volcanic explosions and due to strains set during cooling. The speed of dome growth varies considerably, but some rise by as much as 25 m/day. External Features of Volcanic Domes External Features of Volcanic Domes The exteriors of volcanic domes are distinguished by several distinctive geomorphic features that develop at different stage of growth: (a) Subsidiary Flows: The cooling outer part commonly ruptured by internal stretching, and lava oozes through the ruptures, forming trickles that move for varying distances down the side of the dome. The degree to which a growing dome spreads out from the margin of the vent depends on the magma viscosity. Some domes spread very little, whereas others spread out several times their height, and grade into short thick lava flows. Occasionally, part of the dome may break away, and form a short lava flow. (b) Crumble Breccias: Many of the angular crustal blocks pushed and jostled by the growing dome, and on the edges of the dome, break free and roll down the sides of the dome to come to rest against its base. These accumulation of fallen crustal blocks form banks of loose rock fragments around the edge of the dome, and are known as crumble breccias. These breccias may extend up the sides of the dome giving it a conical shape. The deposits are commonly massive, but sometimes show a suggestion of bedding due to the crude parallel orientation of elongate fragments. There may also be occasional layer of ash, which accumulates on the surface of the growing dome during explosive activity. 32(c) Spines: These structures may form by squeezing out of viscous magma from within ruptures in a solid to semi-solid shell. The largest spine observed historically was the spine that formed during the 1902 eruption of Mt. Pelée, Martinique. This spine, which had a 9 month growth period, reached a height of about 300 m above the top of the summit dome, and was 1000 m in diameter at its base. Without the loss of material due to crumbling, the Peléean spine would have grown to a height of almost 900 m. Spines are commonly angular in cross-section during their early stages of growth, but become more rounded due to wearing away of the aperture through which they are thrust. The great spline of Pelée had a nearly vertical straight face and an opposite face that was characterized by a curved surface that was polished and striated up the upward movement of the spine through the aperture in the solid carapace of the dome. Spines are commonly accompanied in their development by huge, steep-sided pyramidal or conical peaks called pitons. (d) Coronet Explosions: Volcanic explosions commonly occur around the base of domes and spines. The surfaces of separation between domes and the surrounding rocks, or between the spine and dome crust, constitute zones of weakness that allow gases to escape from below and within the dome. Jets of gas may issue from around the base of the dome like spikes around a crown, and thus are called coronet explosions. Explosions at the base of the Chaos Crags dome, Mount Lassen at 300-600 years ago, resulted in collapse of part of the dome, and led to the formation of great avalanches that rushed about 5 km down valley leaving behind a blocky deposit now known as the Chaos Jumbles. Internal Structures of Volcanic Domes Internal Structures of Volcanic Domes Some endogenous domes have a series of concentric, onion-like layers, that result from the gradual expansion from the mass of a somewhat inhomogeneous magma. More commonly, domes are either essentially structureless, or show a divergent, fan-like structure in which ribs of the fan radiate upward from the vent. The fan-like structure is generally shown by aligned phenocrysts and layers of varying composition or by vesicularity drawn out by differential flowage. The fan- structure has a two-dimensional cross-section of nested cones which points downward, and is sometimes expressed as concentric fractures on the dome surface. The internal structures of domes change from horizontal to vertical upward from the vent. 33VI. Products Of Volcanic Explosions VI. Products Of Volcanic Explosions Terminology and Classification Terminology and Classification Fragments which are thrown out by volcanic explosions are referred to collectively as ejecta, and accumulations of these fragments are known as pyroclastic rocks or tephra depending on whether they are consolidated or not. Different sorts of volcanic explosions produce somewhat different types of ejecta. Instantaneous or long-continued, weak or violent, all volcanic explosions are the result of escape of gas from the magma, but they may be subdivided according to the origin of the gaseous component as: (a) if magmatic or juvenile gases, then magmatic explosions; (b) if steam generated where water comes in contact with hot rock or magma, then phreatic explosions; or, (c) if both gas sources, then phreatomagmatic explosions. Ejecta produced by these types of volcanic explosions have been classified by several different criteria: (a) origin; (b) fragment size; (c) condition at time of ejection and at time of striking the ground; and, (d) degree of consolidation of the deposit. Origin Considering origin first, it must be recognized that ejecta may be derived from the molten magma or from rock that was already solid (non-magmatic ejecta). Non-magmatic ejecta may represent: (a) already solidified magma of the same eruption; (b) rocks of the same volcano but formed during earlier eruptions; or, (c) rocks derived from the underlying crust and unrelated to volcanic activity (termed accidental ejecta). Magmatic ejecta and Type I non-magmatic ejecta, which are derived from molten magma of the same eruption, are termed essential ejecta, and are typically partly or entirely glass (vitric). Type II non-magmatic ejecta or fragments of older rocks formed during previous eruptions are termed accessory ejecta, and most are partly or wholly crystalline (lithic). Some accessory ejecta consist of coarse-grained clots of several minerals that represent cognate material that was torn from the conduit walls or from parts of magma crystallized at depth. Fragment Size The most important classification of pyroclastic rocks is based on fragment size, although fragments greater than 5 cm in average diameter are subdivided further on the basis of shape, which reflects their physical condition at the time of ejection: 34(1) Bombs, typically composed of low-viscosity basaltic magma, are ejecta larger than 64 mm in average diameter that are thrown out of the vent in a molten state. Highly to moderately fluid, magma may be ejected both as long, irregular strips or as discrete blebs. Because they are fluid, their shape is typically modified during flight through the air and such fragments are typically termed fusiform:: (a) Strips that break up into short segments form cylindrical or ribbon bombs, which are more less circular or flat in cross-section, and typically show twisted, longitudinal fluting. (b) Large blebs pulled up into spheres by the surface tension of the magma form spherical bombs. (c) Fragments that spin during flight form spindle- or almond-shaped bombs, characterized by longitudinal fluting, and one side smoother and broader than the other. The smooth or "stoss" side represents the front side as the bomb fell through the air, whereas the "lee" side is produced by frictional resistance of the magma dragging the still-plastic skin of the bomb toward this side. this resistance often forms a thin projecting rim along the edge of the stoss side: (d) Bombs of very fluid magma, that is projected only to moderate heights and strikes the ground while still liquid, flatten or even splash to form pancake or cow-dung bombs. At the other end of the spectrum, very viscous bombs are not rounded during flight, and although their outside is nearly solid, the inside is still plastic enough to expand as gases escape and produce a skin that is broken by deep cracks, forming what is called bread-crust bombs. Most bombs are simply an irregular and generally extremely vesicular lumps, which are described as cinder or scoria. Fusiform bombs may only form at the very end of an eruption, because they represent denser material than scoria or cinder, and they form at a stage when the amount of gas in the magma started to decrease. Most bombs in cross-section are at least somewhat vesicular, and they are often characterized by concentric layers of greater and lesser vesicularity. Exceedingly vesicular cinder are called pumice. Pumice of rhyolitic magma are characterized by vesicles that are stretched out into long very thin tubes, giving the fragment a silky appearance. In contrast, far less abundant basaltic pumice typically consists only of thin glass threads that mark the intersections of vesicles. These form the lightest rock (0.3 gm/cm 3 ) known to exist, what is called thread-lace scoria or reticulate. Bombs that have formed around a core of older, solid accessory or accidental fragments, are termed cored bombs. Showers of still-fluid blebs striking ground around the vent may flatten and mod themselves to the underlying surface, forming an accumulation of flattened and welded fragments called spatter or agglutinate. Masses of tephra containing a large proportion of bombs is called agglomerate. Most bombs are less than 25 cm in diameter, but some may be exceptionally large, e.g. 6 m irregular, elongate bombs at Paricutin; up to 1 m fusiform bombs at Mauna Loa; and, up to 1.3 m cow-dung bombs at Stromboli (1965). 35(2) Blocks: These are angular ejecta larger than 64 mm in average diameter that are thrown out of the vent in a solid state. Blocks typically are formed by the disruption of the crust of a lava pool or a dome. They may be entirely cold or still warm and incandescent when deposited. Accumulations of blocks are called breccia, and it is often desirable to specify whether the deposits are pyroclastic breccia or phreatic breccia depending on the type of volcanic explosions responsible. (3) Lapilli: These are ejecta between 2 and 64 mm in average diameter, may be essential, accessory or accidental in origin, and may be ejected in either a liquid or solid state. Lapilli are the most abundant type of fragment in cinder deposits. Special forms characterize drops of basaltic lava that are ejected in a very fluid condition and solidified in the air: (a) droplet-shaped fragments form Pele's Tears, and (b) fragments drawn out into threads form Pele's Hair. An unusual type of lapilli-sized fragment is known as accretionary lapilli, which grow by accretion or addition of concentric layers of fine moist material to a nucleus, like the growth of hailstones. (4) Ash: This is tephra less than 2 mm in average diameter, may also be essential, accessory or accidental in origin, and ejected in either a liquid or solid state. Depending on the material that composes the ash, it can be described as: (a) lithic, composed dominantly of solid rock; (b) vitric, composed dominantly of glass; or, (c) crystal, composed dominantly of crystals. The most common type of ash is vitric. Unconsolidated deposits of ash-sized material are referred to as ash layers, ash beds or ash blankets. Consolidated ash deposits are described as tuffs, or more specifically, depending on the predominant constituent as: (a) lithic tuff; (b) vitric tuff; or (c) crystal tuff. Some ash deposits contain moderately to very abundant lapilli, blocks or bombs, and are termed lapillistone, lapilli tuff, tuff breccia or tuff agglomerate. Airfall Ash Deposits Airfall Ash Deposits Tephra produced by a volcanic eruption may be distributed by fall through the atmosphere or by flow over the ground surface. Tephra may also be dispersed by ocean currents, where ash has fallen on seawater and coagulated. The most widespread pyroclastic product is airfall ash deposits. Tephra of basaltic eruptions are much less voluminous than those of intermediate to rhyolitic eruptions due to the less explosive style of basaltic volcanic activity. 36Dispersal An eruption column can carry ash-sized fragments to altitudes of 6-50 km above the vent. The dispersal of ash from the eruption column depends largely on the directions of winds at intermediate and high altitudes (4500-13000 m). At high levels, atmospheric flow is laminar, but at low levels, it is turbulent. Ash can be transported at 100-200 km/hr in the upper atmosphere. Once particles move into the upper atmosphere, however, velocity decreases due to: (a) gravity; and, (b) air resistance. The rate of ash falling from the highest point of its trajectory increases until the acceleration of gravity is balanced by the decelerating effect of air resistance. Beyond that point, the velocity remains constant, and ash can remained suspended until the wind velocity drops below the particle's settling velocity, V t , defined by Stokes Law: V t = {(8Rr s g)/(3Cr v )} 0.5 where R is the particle radius, r s is the particle density, C is the drag coefficient of the particle, and r v is the density of the transmitting medium. In general, greater amounts of tephra fall out of the ash cloud near the vent, so that airfall deposits typically thin away from the vent. However, secondary thickness maxima may occur downwind. Airfall deposits typically have a circular or regular to irregular, fan-shaped distribution with respect to their source. The azimuth of the fan axis may change with distance from the source, and thickness may be skewed to one side, perpendicular to the fan axis. Moreover, the apex of the dispersal fan may no be on the volcano, such as at Mount St. Helens or White River, Yukon. Structures Of Airfall Deposits Because they form as atmospheric fallout, these deposits are characterized by what is termed mantle bedding, as they typically "mantle" or drape over the underlying topography except where it is rugged. Bedding planes are distinct where deposition is on weathered or erosional surfaces, or different rock types. They may be gradational if deposition is slow by small increments so that bioturbation, wind reworking, and other soil-forming processes dominate. The fabric in beds is commonly isotropic because elongate fragments are uncommon, with the exception of platy minerals and glass shards. Airfall deposits are generally well-bedded and well-sorted, with bedding becoming more pronounced as sorting increases, and with size and sorting parameters varying geometrically with distance from the source within single layers. Inman parameters (s f ) are commonly 1.0 to 2.0 within both relatively coarse-grained as well as fine-grained tephra. Median particle diameters (Md f ) are commonly -1.0 to -3.0 (2 to 8 mm) or smaller (phi values) close to the source, but farther away may vary from 0.0 (1 mm) to 3.0 (1/8 mm) or more The sorting of airfall deposits depends on: (a) the distance from the vent; (b) variations in the strength and duration of eruptions; 37(c) length of quiescence between explosions; (d) changes in the direction of fragment ejection; and, (e) the direction and velocity of the wind. Within products of a single eruption, bedding tends to show normal grading, but reverse grading may occur in waterlain pumiceous deposits, and cross-bedding may result due to shifts in wind strength and direction. Differences in the proportions and densities of lithic, crystal and vitric constituents produce both lateral and vertical variations in the size and nature of particles and nature of the deposits. Vertical variations usually show increasingly basic compositions, often reflecting the tapping of compositionally zoned magma. Lateral variations reflect differences in the settling velocities of particles. Most ash deposits become more silica-rich with distance from the vent, as different minerals are winnowed from the ash cloud. Morphology of Ash Particles Ashes are best placed into two broad genetic categories: magmatic and phreatomagmatic. Ashes from magmatic eruptions are formed when expanding gases in the magma form a froth that loses its coherence as it approaches the ground surface. During phreatomagmatic eruptions, the magma is chilled and fractured on contact with ground or surface waters, resulting in violent steam eruptions. In low-viscosity magmas droplet shape is, in part controlled by surface tension, by acceleration of the droplets after they leave the vent, and by air friction. The ash particles consist of mostly sideromelane, translucent basaltic glass, or tachylyte, opaque Fe-Ti oxide charged glass. The sideromelane particles exhibit smooth, fluidal surfaces and a thin skin. In higher viscosity magmas, the morphology of ash particles is controlled primarily by vesicle density and shape, the vitric fraction generally consisting of very angular pumice fragments and thin vesicle walls broken from pumice fragments during or after eruption. The morphology of lithic fragments is dependent on the texture and fracture pattern of the rock type broken up during the eruption. The morphology of ash particles from phreatomagmatic eruptions is controlled by the thermal stresses within the chilled magma, which result in fragmentation of the glass to small blocky or pyramidal glass particles. Vesicle density and shape play a minor role in determining the morphology of phreatomagmatic ash particles. Pyroclastic Flow And Surge Deposits Pyroclastic Flow And Surge Deposits Pyroclastic flow and surge deposits represent the movement of large volumes of pyroclastic material with the general behavior of lava flows. They act as heavy fluids controlled in their movement by gravity and the topography of the underlying land surface. The flows may be related to dome collapse, to explosive activity at the crater of composite volcanoes, or to discharge from fissures. The deposits, which have variously been termed ashflow tuffs or ignimbrites, are diverse and reflect different types of eruptions and depositional regimes. From the perspective of flow mechanic, there are essentially two kinds of deposits: (1) Pyroclastic flow deposits which are commonly poorly sorted and massive, and, (2) Pyroclastic surge deposits which are better sorted, finer-grained, thinner and better bedded than pyroclastic flow deposits. 38Pyroclastic flows originate in different tectonic and volcanic settings and have vastly different volumes. Eruptions producing pyroclastic flow deposits on the order of 0.001 to 1.0 km 3 are from small central vent volcanoes typical, but not confined to magmatic arcs: (a) Mount Pelée (b) 1968 Mount Mayon, Philippines (c) 1976 Augustine volcano, Alaska (d) 1980 Mount St. Helens, Washington (e) 1982 El Chichon, Mexico More voluminous flows of 1-100 km 3 originate from larger stratovolcanoes: (a) 1883 Krakatoa, Java (b) Mount Mazama Volumes of 100-1000 km 3 are associated with the formation of large caldera which develop as a consequence of eruption of such large volumes and not necessarily at the site of a pre-existing volcano: (a) Yellowstone, Wyoming (b) Valles, New Mexico (c) Long Valley, California (d) Silverton-Creede, Colorado (e) Toba, Sumatra (f) La Garita, San Juan Mountains, Colorado ( where a single mapped pyroclastic flow sheet is greater than 3000 km 3 ). In general, small- to intermediate-volume flows range from rhyolitic to basaltic in composition, whereas large-volume flows are most commonly rhyolitic to dacitic. Each type of pyroclastic flow consists of different types of fragments that reflect how the flows originate: (a) Small-volume flows produced by dome collapse or explosions associated with dome formation commonly contain abundant poorly vesiculated products of the domes, although some dominantly pumice flows also occur. (b) Intermediate- to large-volume flows are usually composed entirely of highly vesiculated materials derived from the rapid vesiculation of magma. Relationship to Topography Pyroclastic flows may completely drain from upper slopes and only be preserved in the lower parts of valleys, thereby becoming initially thicker away from the source. Pyroclastic flows may be confined to valleys. On the upper slopes of volcanoes, pyroclastic flows drain down valley centers, leaving levees or high-water marks and larger rock fragments on both sides of a valley, or along the outer edge of a sinuous channel because of the momentum of flow. Beyond mountain slopes, pyroclastic flows spread out in fan-like lobes. Widespread sheet-like ignimbrite layers associated with large calderas and other volcano-tectonic depressions commonly have sufficient volume and thickness to smooth out underlying topography. They: (a) thicken and thin according to topographic irregularities underneath; 39(b) maintain a nearly flat, horizontal or gently sloping surface; and, (c) gradually thin toward their distal edges. Successive flows of great volume may mask completely previous topography. Pyroclastic surges can spread over topography of moderate relief and override the sides of a valley, and their deposits may mantle topography similar to fallout tephra, but unlike fallout tephra, they can become ponded and thin toward valley margins. Flow Units and Cooling Units A basic stratigraphic and field distinction that must be made for intermediate- to large- volume pyroclastic flow deposits is the difference between flow units and cooling units. A flow unit is a depositional unit that represents a single pyroclastic flow deposited in one lobe. The thickness of individual flow units can vary from a few centimeters to tens of meters, and the lobes may follow one another within minutes or hours. The boundaries between flow units are marked by: (a) changes in grain size, (b) composition, (c) fabric, (d) concentration of pumice lapilli or block accumulations, and, (e) cross-bedded zones. When several very hot flow units pile rapidly one on top of the other, they may cool as a single cooling unit. A simple cooling unit forms when a single flow or successive flows cool as a unit with no sharp changes in the temperature gradient. A compound cooling unit forms when there is an interruption in temperature that disturbs the continuous cooling unit zonation of successive hot flows. Cooling from emplacement- to ambient temperature may take tens of years, depending on the thickness of the deposit and the emplacement temperature. Thus, many ash flow deposits mare mapped as cooling units, even though they composed of several flow units. A cooling unit is marked by a more or less symmetric patterns of zones of rock differing in degree of welding and thus density resulting from different cooling regimes. In young deposits: (a) The top and bottom parts of cooling units are commonly composed of friable unwelded pyroclastic material. The basal layer is unwelded because it cools quickly against the cold rock basement, and the top because of relatively rapid heat conduction and radiation into the atmosphere. (b) The area of densest welding, which occurs in the lower half of a cooling unit, is that zone that remains longest at the maximum emplacement temperature. At high emplacement temperatures and slow cooling rate, partial or complete crystallization (primary devitrification) of hot and compacting glassy pyroclasts occurs in the interior of thicker cooling units. Such zones grade into poorly welded zones lithified by crystallization of high temperature vapor-phase crystals, typically silica polymorphs (tridymite, cristobalite) and alkali-feldspar. 40(c) The most obvious breaks between cooling units is the presence of an erosional surface, although breaks may also be indicated by reversals in porosity and flattening ratios of pumice fragments. Components Pyroclastic flow and surge deposits are composed of crystals, glass shards and pumice, and lithic fragments in highly variable proportion, depending upon the composition of the magma and the origin of the flows. Ash-flow tuff, by definition, is composed of more than 50 percent of components in the ash size range (<2 mm). These particles form a matrix into which varying amounts of pumice lapilli or lithic lapilli are present. (a) Glass Shards The most common ash-sized clasts are glass shards, usually accompanied by smaller amounts of pumice particles. The ash and lapilli- sized pumice fragments are characterized by either spheroidal or long subparallel tubular vesicles a few millimeters to micrometers in diameter. (b) Crystals These are the next most common ash-size component. Crystals in ignimbrites, which range from 0 to 50 percent abundances, are commonly broken. Phenocrysts within accompanying comagmatic pumice lapilli or blocks, however, are largely non-fragmented, which indicates that breakage occurs during eruption and transportation. Breakage may continue even during compaction. Crystals are generally more abundant in the matrix than in pumice lapilli and bombs, suggesting preferential concentration in the matrix relative to glass shards during transport. (c) Lithic Fragments The lithics rarely exceed 5 volume percent of intermediate- to large-volume and some small-volume pumiceous pyroclastic flows. There are three major sources for lithic fragments: (a)slowly cooled and crystallized magma rinds from chamber margins; (b)rocks from conduit walls; and, (c) rock fragments picked up along the path of the pyroclastic flow. Characteristics of Ash-Flow Deposits Characteristics of Ash-Flow Deposits Most unwelded pyroclastic flow deposits are poorly sorted and massive, but may show subtle grading, alignment bedding or imbrication of oriented particles. In contrast, most pyroclastic surge deposits are thinner, finer-grained and better sorted than flow deposits, and wavy- or cross- bedded structures may be common. Internal Layering Many features, including graded bedding, give evidence of emplacement as high- concentration laminar flows. Layering is manifest within pyroclastic flow deposits as: (a) graded basal zones, (b) discontinuous trains of large fragments, (c) alternating coarse- to fine layers, 41(d) crude orientation of elongate or platy particles, and, (e) color or compositional changes. Grading within a single flow unit can be normal, inverse symmetrical, or multiple. Pumice fragments may be inversely size graded and lithic fragments normally graded within the same horizon because of their widely different densities. Slight differences in size of fragments in different layers may also give an irregular and indistinct stratification to the deposit. Flat fragments within pyroclastic flow deposits may be strongly oriented parallel to depositional surfaces , or may become imbricated near the basal parts of the deposits due to shearing forces within the flow during laminar movement. Multiple grading may result from: (a) graded basal zones, (b) a continuous recurrent surging within a single flow; (c) mechanical differentiation due to shearing during laminar motion within a moving flow; and, (d) separate flows repeated at relatively short time intervals. Gas-Escape Structures The welded portions of pyroclastic flow deposits may be characterized by large vapor pores or cavities called lithophysae. Pyroclastic flow deposits are also commonly characterized by numerous degassing pipes or fumaroles. These degassing pipes may be recognized a oxidized zones or as fines-depleted pipes in lower-temperature deposits. Minerals precipitated by fumaroles may built fumarolic mounds and ridges, and are numerous where crystallization of an ash sheet is most intense; the vapor-phase deposits involve minerals such as tridymite, alkali-feldspar, hematite and sulphate minerals. They are generally absent where a sheet is thick, densely welded and vitric. The mounds, such as those of the Bishop Tuff in the Long Valley Caldera, California, may be relatively large features that stand 0.5 to 15 m above the flow surface and are up to 60 m in diameter; the ridges may be as long as 600 m. These zones typically represent zones of intense devitrification, the late stages of which may lead to development of spherulites. Textural Relationships Most pyroclastic flow deposits are poorly sorted, having sorting values (s f ) that are greater than 2.0, and tend to decrease, as do median diameter values (Md f ) with length of transport. Size distribution curves of different types of pyroclastic flow deposits tend to follow a Gaussian distribution, which indicates that sorting takes place during eruption and transport. Poor sorting is characteristic throughout the length of a single pyroclastic flow sheet, but sorting varies vertically at any single locality and tends to improve slightly with distance. Maximum size of both lithics and pumice decrease with distance in subaerial deposits. In the textural analysis of pyroclastic flow deposits, it is important to know the relative proportions of pumice, lithics, and crystals because their size distributions, sorting and other parameters of these three subpopulations may differ for reasons other than sorting in the eruption column and during flow. Glass particles may diminish is size toward the margins and distal ends of the flow, due to breakage during transport. Lithic fragments could be derived from magmatic stoping, by fragmentation of the walls of the magma chamber and vent, or fragmentation of a plug dome or dome within the vent, and they may also be picked up from the ground during flowage. The size distribution of crystal fragments is a function of original phenocryst sizes in the magma and of breakage during explosive eruptions. Pumice has low mechanical strength and therefore 42may be reduced in size during eruption and flow, causing a preponderance of pumice dust in the fine-grained fraction of the deposit. Segregation of Crystals and Lithics Segregation and eltruiation of particles occur within the conduit and the eruption column, and during flowage, causing enrichment of crystals and lithics, and depletion of fine-grained vitric particles. Fragments with relatively low settling velocities are carried high into the atmosphere and do not enter the flow, and others escape from the top of the flow while it is moving. These processes give rise to a distinctive type of fine-grained fallout tephra that has a crystal/vitric ratio symmetrically lower than that from manually crushed pumice. Crystal enrichment may be greatest in the basal layers than in the middle of some ignimbrites, due to selective loss of glass shards from the flow and to a reduction in pumice size by abrasion during flow. Temperature Effects Pyroclastic flows, although they are particulate material, are a good heat-conserving mechanism. Mixing of hot pyroclastic material with cold air during flowage is minimal and restricted to a thin surface of the flow. Hot pyroclastic flows may be nearly at magmatic temperatures during movement and shortly after emplacement. The emplacement temperature is determined by: (a) the liquidus temperature of the magma, (b) height to which material is lifted in the eruptive column, and, (c) total volume of a flow. Temperatures measured within hours to days, to depths of several centimeters to meters within observed pyroclastic flow deposits, generally range from about 500° to 650°C. Mount St. Helens emplacement temperatures ranged from 750° to 850°C near the vent, and 300° to 730°C farther away, with latter eruptions being hotter than earlier ones, and temperatures within individual flow deposits not decreasing substantially along their flow paths. The cooling rate is initially rapid on the surface, followed by a slow decline towards the center. Welding and Compaction High emplacement temperatures are responsible for some of the most characteristic features of ignimbrites, e.g. the plastic deformation and welding together of glass shards. In compaction of an ashflow deposit, two kinds of compaction are present: (a) mechanical compaction: Mechanical compaction takes place as the result of simple loading without significant change in particle shape. Particles maintain their relative positions after coming to rest except that elongate particles tend to be rotated toward the horizontal. Pumice fragments generally maintain a random orientation. Thus, the primary effect of mechanical compaction is to produce reduced porosity: (b) welding compaction: Welding compaction results from variable viscous deformation of vitric fragments, from completely undeformed shard typical of low-temperature fallout deposits to nearly homogeneous solid glass typical of obsidian in which there are only ghost-like outlines of former shards in a continuous glass matrix. The main control is the amount of time that 43temperatures remain above the threshold for welding (about 550°C). The threshold temperature depends on volatile content and chemical composition of the glass. Complete welding may cause collapse of porous pumice to form dense, black, glassy discs, or fiamme, which are normally drawn out in a parallel structure or banding referred to as eutaxitic texture. The degree of welding is also dependent on load pressure, but less so than temperature, viscosity and volatile content. Ashflow tuffs tend to be denser toward their source, where welding is possible due to higher temperatures despite reduced thicknesses. Columnar jointing is common in moderately to highly welded tuffs. Structures Related to Temperature and Viscosity Several structural features of welded to partly welded tuffs are related to the effects of viscosity that develop during movement and deflation of a pyroclastic flow. Three main groups of features are especially useful for determining flow direction: (1) those induced during inflated movement; (2) structures caused by dense, lava-like flow or creep during deflation after emplacement and shortly before coming to complete rest; and, (3) those caused by compaction after the flow has ceased forward movement. These flowage structures include: (1) zones of densest welding and maximum stretching of pumice; (2) stretched and pulled apart pumice with (a) cracks convex toward the flow front, or (b) broken and rotated segments with rotation toward the flow direction; (3) tension fractures in welded matrix adjacent to unbroken pumice clasts with cracks dipping in the direction of movement; (4) spindle shape structure around rotated inclusions and strongly developed imbrication; (5) folds with axial planes that dip sourceward; (6) imbricated stretched pumice dipping up-flow; and, (7) ramp structures showing asymmetry. Flow foliations are also manifested by flattened and elongate gas cavities. Post-depositional compaction of high-temperature pyroclastic flows causes flattening and welding of vitric shards and pumice. Measurements of flattening ratio (F = length/width) of pumice fragments have revealed a steady increase in the flattening from the top downward into the body of ash-flow sheets despite generally parallel changes in density. Classification and Nomenclature of Pyroclastic Flows Classification and Nomenclature of Pyroclastic Flows Classification and nomenclature of pyroclastic flows and their deposits have been the subject of much confusion and debate. The wide range in physical properties of the erupting magmas, the different ways that flows originate, differences induced by transport processes, and differences in compaction and cooling structures and textures have given rise to a large array of 44names for the flows and their deposits. Traditionally, different kinds of pyroclastic flow have been named according to the volcano where first observed. Williams (1957), for example, arranged the different types of flows in order of increasing gas content, increasing volume, and diminishing viscosity of the initiating magmas: 1. Flows related to domes or to crumbling fronts of lava flows: (a) Merapi type. Flows form by non-explosive disintegration and collapse of the flanks of domes and summit-spines or by breakup of the snouts and levees of viscous lava flows on steep slopes. These flows result gravity- collapse of the oversteepened flanks of domes and unstable summit- spines. They may be triggered by earthquakes or by internal expansion of domes. They are not buoyed up by exsolving gases, but rather slide. The chaotic, unsorted, and unstratified deposits are generally less extensive than those caused by explosions; they differ also in containing less pumice, and correspondingly more angular blocks. (b) Peléean type. Flows form by explosive eruptions immediately before or during the rise of domes. These are produced by explosions before and during the rise of volcanic domes. Some Peléean flows are caused by low-angle blasts. The ejecta vary from almost wholly lithic to almost wholly pumiceous. Largest, most gas-rich, and most destructive avalanches usually occur during the initial phases of the growth of domes; these early flows are composed entirely of fresh effervescing magma in the form of ash; blocky lithic debris is quite subordinate. Subsequent avalanches issue from the flanks of growing domes, and the deposits being much coarser and include abundant lithic blocks. Deposits vary greatly in form and texture, are almost wholly confined to topographic depressions. They are heterogeneous, unstratified and unsorted. They are made up of angular blocks, torn from the solid carapace of the parent dome, subangular, still effervescing bombs from the plastic interior, and sand- to dust-sized debris, some derived from effervescing magma and bursting bombs and some from larger fragments broken in transit. Some blocks are of enormous size: up to 8 by 5 meters across. Almost all fragments, particularly those in which the proportion of fresh magma was large at the time of eruption, are characterized by a high degree of porosity. Layers and lenses of sand- and dust-sized ash are commonly interbedded with the coarser debris. Peléean deposits are characteristically monolithologic and consist wholly of accessory and juvenile debris of uniform composition. Accidental lithic fragments are rare. The deposits are not welded, though they may be indurated by compaction. 2. Flows from summit-craters: (a) St. Vincent type. Flows produced by backfall of ejecta from the margins of vertical eruption-columns. The pyroclastic flows originate by "backfall" of the outer, more slowly rising parts of the eruption-column, produced by gravitational collapse. These columns consist of two parts: A dense lower part first accelerates by decompression of the gas and then accelerates as it interacts with the atmosphere; a lighter upper part rises 45because it has a higher temperature and hence a lower density than the atmosphere. When the fall-out of ejecta and heating of entrapped air fail to reduce the effective density of the eruption-column below that of the atmosphere, the mixture of gas and solids falls back toward the vent and spreads outward as a glowing avalanche. Reduction of the gas-content of an erupting magma may produce a change from pumice showers to pumice flows. Most, though not all, flows are preceded by pumice falls, and the size of ejected fragments often increases as the eruption proceeds. The great mobility of pyroclastic flows of the Saint Vincent type is explainable by the mechanism of collapse of the mixture of hot gas and solids; the higher particles reach in the eruption column, the greater the distance of pyroclastic flow after collapse. The deposits are unstratified or poorly stratified, poorly sorted, and unwelded, and are characterized by their fine texture and richness in crystals. Bombs and blocks make up only 3 to 5 percent of the total volume; even lapilli constitute only a small proportion. More than 90 percent of the deposits are of sand-size. Crystals make up 45 percent of most bombs, but they make up 73 percent of the ash. The concentration of crystals relative to vitric ash in deposits only a few hundred meters from the crater rim shows that gravity- separation takes place in the eruption-column. (a) Krakatoan type. Pumiceous flows discharged by voluminous upwelling from summit-vents and circumferential fractures on composite cones; eruptions commonly result in the formation of a caldera. Pumiceous pyroclastic flows erupted from summit-vents of large composite volcanoes during late stages in their history. They are almost invariably preceded by pumice falls, and usually involve magmas that are more fluid than those of the Peléean, St. Vincent, or Asama types. Consequently, there is much more pumiceous material among the deposits and correspondingly less lithic debris. Pumice flows follow pumice falls as gas pressure diminishes. The pumice falls are well-sorted, well-stratified, and diminish in thickness away from the source. The deposits of the pumice flows are poorly sorted, and their stratification is irregular and indistinct. Except near the top, where coarse fragments are concentrated, they consist chiefly of sand-sized pumice. The compositions of the pumice-fall and pumice-flow deposits are essentially identical. The surfaces of the flows, particularly near their edges, are characterized by subparallel ridges, up to a meter high, composed of large pumice lumps that were segregated from the finer pumice by differential rates of flow. The flows thicken toward their terminus. (b) Asama type. Flow formation intermediate between those of Peléean type and those of St. Vincent and Krakatoan types. Initial discharge of gas- rich magma produces a pumice fall followed by two pyroclastic flows of diminishing gas-content. The pyroclastic flows, which display features transitional between those of Peléean and Krakatoan flows, are issued from a summit crater as the magma foams over the crater rim and sweeps downslope. The vesicularity of the ejecta decreases during the eruption. The earliest flows consist of material plastic enough to anneal; even at the margins of the flow, where the deposits are less than 30 cm thick, they are compact and crudely jointed. No fragments, however, show 46flattening. Later pyroclastic flow involve more viscous magma, and many of the ejected fragments are solid; these flows do not spread as widely as the early flows, and tend to be confined to narrow channels. The deposits consist of dense or slightly vesicular blocks and bombs and little juvenile ash. Some blocks measured up to 30 m across None of the later flows are welded or indurated by compaction. 3. Flows discharged from fissures: (a) Valley of Ten Thousand Smokes type. Eruptions from one or more short linear fissures, the location and trend of which are unrelated to a cone or crater. These flows consist predominantly of sand- and dust-size pumiceous particles mingled with lapilli, and are almost completely unstratified. Bombs and lithic blocks are rare. The deposits are almost completely unsorted. Lenses of cross-bedded, fluviatile pumice separate some of the flows; these are laid down by floods when rivers burst dams. Most deposits are only weakly indurated, but some are slightly welded and show columnar jointing. The flow are not characterized by distortion of glass shards or flattening of pumice lumps. (b) Valles type. Eruptions of siliceous pumice from arcuate fissures formed by regional arching of the roofs of large bodies of rising magma; volumes of ejecta are usually so great that the roofs of the magma chambers collapse along the arcuate fissures to produce calderas. These voluminous pyroclastic flows are discharged from arcuate fissures that develop where either there has been little or no volcanism for long periods or where long-continued eruptions have produced thick volcanic deposits. All originate above large bodies of siliceous magma that dome their roofs. Once the fissures open, foaming magma escapes in huge volumes, first producing airfall deposits, and then ignimbrites. The volume of air-fall pumice is much less than that of the ignimbrites and the latter are commonly welded. So much magma is expelled that the roof of the magma chamber usually collapses along the arcuate fissures. Alternatively, Wright and others (1980) named the flows and their deposits according to field criteria such as relative amounts of poorly vesiculated blocks, and ash (see Tables 12-1 to 12-6 in Cas and Wright, Volcanic Successions). They emphasized (1) vesiculation of essential fragments (related to gas content, viscosity, and rate of gas release), and (2) eruptive mechanisms. 47VII. Laharic Deposits VII. Laharic Deposits General Features General Features Lahar is Indonesian for volcanic breccia transported by water, synonymous volcanic debris flow, a mass of flowing volcanic debris intimately mixed with water. The term refers both to the flowing debris-water mixture, and the deposit thus formed. Many lahars are associated with stratovolcanoes of which they may comprise significant volumes of bulk. Most lahars are relatively limited in extent, and occur in valleys or on alluvial aprons or lowland areas immediately surrounding volcanoes. Many lahars are initiated directly by volcanic eruption, whereas others originate in ways similar to nonvolcanic debris flows. Once flow begins their fluid characteristics appear to be similar or identical. The flows are non-Newtonian fluids that have a yield strength, behaving like plastic material similar to wet concrete, have a high bulk density, and exhibit the property of strength which greatly influences the final textures and structures of the deposit. The Newtonian properties of water (i.e. lacking in yield strength) begin to be modified by particle interference when the volume of solids exceeds 9 percent. At volume concentrations of about 20 or 30 percent, particle interactions almost completely dominate flow behavior. The flows are fluids in which the water and solids form an intimate mixture that flow with laminar motion. As velocity decreases, the entire flow stops rather abruptly, after which water separates from the granular material by percolation or evaporation. On steep slopes, velocities may be rapid enough to keep the entire mass in motion, but as slope decreases, the mass congeals unless it is thick enough to maintain a high shear stress at the base of the flow. As the gradient decreases, velocities decrease, and the flow thins, shear stresses increase until the flow congeals to its very base and deposition is complete. Lahars follow pre-existing valleys and may be interstratified with alluvium, colluvium, pyroclastic rocks of diverse origin and lava flows derived from the same source area. They leave thin deposits on steep slopes and in the headwaters of valleys, but become thicker in valley bottoms and form fans that coalesce or else form broad individual lobes in lowland areas on very low slopes. Movement of lahars down valleys generally occurs in surges, or peaks of flow. During their course down a valley, lahars tend to leave thin "high water" marks (veneers) where a constriction momentarily causes a large debris flow to pond up to several tens of meters above the valley bottom and then drain away. Lahars vary greatly in thickness. They tend to maintain a relatively constant average thickness on relatively low slopes but locally vary depending on the configuration of underlying topography. Lahars come to rest with steep sloping lobate fronts. Most lahars are probably less than 5 m thick, but some are more than 200 m thick. Surface of Lahars Lahar surfaces tend to be remarkably flat over wide areas but contain local swells and depressions interpreted to be caused by differential compaction over an irregular underlying surface. The form, shape, and size of irregularities depends on the viscous properties of the flows and the number and characteristics of multiple lobes. 48Basal Contact of Lahars Although lahars may be very thick and carry large boulders, they commonly do not erode the surfaces on which they flow except on very steep slopes. Lahars can pick up loose materials from surfaces on steep slopes or where local turbulence develops within the flow owing to highly irregular channels. Some Pleistocene lahars in the southern part of the Puget Sound lowland have traveled 60 to 80 km from their source without picking up appreciable debris from the surface on which they flowed. Components of Lahars Depending upon their origin, lahars may be monolithologic or heterolithologic. Monolithologic varieties are derived directly by eruption, whereas collapse of crater walls or avalanching of rain-soaked debris covering steep volcanic slopes give rise to heterolithologic types. Lahars characteristically contain dense angular to subangular rock of dominantly andesitic to dacitic composition mixed with ash-sized minerals and lithic particles. Many lahar deposits contain charred wood, indicating that they were initiated as hot pyroclastic flows then cooled down during transport. Grain-Size Distribution Particles carried by lahars range from clay- to boulder-size, but the percentages of each size fraction vary enormously from deposit to deposit and within a single deposit. In general, lahars are coarser-grained and more poorly sorted than pyroclastic flow deposits. Grain-size parameters show the obvious fact that lahars have a wide range in grain size and are coarse-grained and poorly sorted. The presence of large boulders, commonly exceeding l m in diameter, is one of the most characteristic features of lahars except perhaps, in their terminal zone. Large fragments progressively decreased in number and size away from the source, although the finer constituent (matrix) may not show corresponding changes. Erratic fluctuations in median diameter are attributed to the longitudinal inhomogeneity of the flow caused by deposition from individual debris tongues that differed in grain size. Grading Many lahar deposits show a subtle grading of the coarse-grained (>2 mm) dispersed phase, but it may not be evident in the matrix phase. Single depositional units generally have an irregular but slightly more concentrated arrangement of large fragments a short distance above the base; such layers are reversely graded. The large fragments in a lahar rarely rest directly upon the depositional surface. Some workers have suggested that large boulders re suspended by turbulence, it has generally been shown convincingly shows that debris flows move in laminar fashion; therefore, large boulders are suspended by combination of high density (buoyancy) and high strength of the matrix. Differences in grading, whether it be absent, weakly or strongly developed, normal or reverse, appear to be related to the relative concentration of solids and fluids; the lower the concentration of solids, the more likely normal grading develops because viscosity, density, and strength of the fluid are less able to support large dense particles as velocity decreases. Where concentration values and viscosity, density, and strength are high, reverse grading develops, especially if the density of fragments is relatively low. 49Fabric The fabric of lahars is commonly regarded as isotropic, but in some lahars disc-shaped pebbles and uncharred twigs and tree trunks concentrated low in the central parts of the deposit are oriented subparallel to base. The development or lack of clast fabric in the flows depends on the mechanism of movement and deposition. Matrix strength in the flows may produce a rigid plug where shear stress is below the yield threshold, and this plug rides on a zone of laminar flow within which the shear stress is greater than the yield threshold. Flow stops when the plug expands to the base of the flow at the expense of the zone of laminar flow Origin of Lahars Origin of Lahars The mechanisms of lahar formation can be grouped into three major categories: l. Direct and immediate result of eruptions: eruptions through lakes, snow or ice; heavy rains falling during or immediately after an eruption; flowage of pyroclastic flows into rivers, or onto snow or ice. 2. Indirectly related to an eruption or shortly after an eruption: triggering of lahars by earthquake or expansion of a volcano causing the rapid drainage of lakes or the avalanching of loose debris or altered rock. 3. Not related in any way to contemporaneous volcanic activity : mobilization of loose tephra by heavy rain or meltwater; collapse of unstable slopes; bursting of dams due to overloading; sudden collapse of frozen ground during the spring thaw. 50VIII. Structures Built Around Volcanic Vents VIII. Structures Built Around Volcanic Vents The accumulation of erupted material around a vent forms a hill or mound. The form of this structure depends partly on the form of the vent and partly on the violence of eruptions. It also depends partly on: (1) the angle of repose of loose fragments; (2) the degree of welding of the fragments; (3) the volume of lava outpourings; and, (4) the viscosity of the magma. The mound may consist wholly of tephra, lava, or some mixture of the two materials. Those built by single eruptions may vary from a few meters to several hundred meters in height, and a few meters to more than a kilometer in width. Repeated eruptions build larger structures. Cinder Cones Cinder Cones Cinder cones are built by "lava fountains" and by eruptions of scoriaceous, basaltic ejecta. They range up to 700 m or more in height, but most are between 30 and 30 m high. Pumice cones resemble cinder cones in form, internal structure, and size range, but they differ in that they are composed of light-colored, more siliceous ejecta, such as rhyolite or dacite. Cinder cones typically represent rapid growth. For example, during its first 24 hours, Paricutin (Mexico) rose between 30 and 50 m; when it was only a week old, it was 140 m high. After a year, it was 325 m high and had a volume of approximately 0.2 km 3 . Over 90 percent of the fragmental ejecta were discharged during the first year. External Form Most cones are essentially symmetrical with slopes of 25° to 40°, and saucer-, bowl-, or funnel-shaped summit-craters. They may overlap or coalesce or even grade into ramparts along a common fissure. The forms of craters are a function of the nature of explosive eruptions and the mechanical properties of ejecta. In general, a shallow explosion focus provide a broad, shallow crater. The initial form is normally funnel-shaped, but slumping of walls, subsidence of material into the vent, or flooding of the crater with lava modifies the shape of the walls and floor. Eruptions of constant strength tend to produce uniform slopes corresponding to the maximum angle of repose of the ejecta. If most eruptions are weak, the slopes of the cones tend to be concave; convex slopes may develop when violent eruptions are more numerous than weak ones. Almost all young subaerial cones less than 1,000 m high and built predominantly of pyroclastic materials have straight or even convex slopes. Larger cones, particularly eroded ones, have concave slopes. Asymmetrical cones and crater-rims result from such factors as: (1) greater accumulation of ejecta on the leeward sides, (2) eruptions from inclined or multiple conduits, or (3) shifting of vents. 51Breached cones, as crescent or horseshoe-shaped remnants, may be caused by lateral explosions; they result usually as lava escapes from vents on the flanks and slices from the upper slopes are rafted away. Internal Structure Ejecta that make up Cinder cones are well stratified, with layers of different coarseness ranging from bombs to ash and consisting mostly of lapilli. Frequent and repeated size changes reflect variations in the strength of the explosions. The ejecta display little gravity sorting; only at greater distances from the vents does tephra show graded bedding. Lava may be interbedded with the scoria, but it is more typical for lava to issue from a vent (bocca) at or close to the base. If present, dikes are irregular in trend and width and may intersect in a complex manner. Most ejecta are coarsest and thickest close to the vent. Ejecta from violent eruptions diminish outward in thickness and size less rapidly than do the ejecta laid down by weak eruptions. Maar Volcanoes Maar Volcanoes Maar volcanoes are low volcanic cones with bowl-shaped craters that are wide relative to rim height. They were originally recognized as small subcircular crater lakes, the term being derived from the Latin "mare" for sea. The various kinds of maar volcanoes are: 1. Maar (sensu stricto): Volcanic crater cut into country rock below general ground level and possessing a low rim composed of coarse- to fine-grained tephra. They range from about 100 to 3000 m wide, about 10 to more than 50 m deep and have a rim height of from a few meters to nearly l00 m above general ground level. 2. Tuff ring: Large volcanic crater at or above general ground level surround by a rim of pyroclastic debris (tuff or lapilli tuff), similar in diameter to maars. 3. Tuff cone: These cones have higher rims, attaining heights of up to 300 m, and are essentially tuff rings where volcanic activity was of longer duration. The distinction between tuff cones and tuff rings becomes arbitrary where one side of a crater stands high and another side low. Most maars result from hydroclastic eruptions; wide craters develop from shallow explosions, subsidence or a combination of both. In groups of nearly synchronous eruptive centers, those erupting on high ground form spatter or cinder cone whereas associated eruption centers in valleys, depressions, or alluvial gravels in coastal regions form maars, tuff rings or tuff cones. Juvenile clasts within the deposits are glassy, non-vesiculated, and have blocky shapes, suggesting that magma was quenched prior to exsolution of volatiles, that breakage of glass resulted from thermal shock and (steam) explosions, and that the vapor and steam phase in the eruption column was partly or largely vapor from external water. Tuff cones and tuff rings are distinct landforms that result from slightly different types of hydroclastic activity an represent a "continuum" of landforms from cinder cones to pillow lava related to environments of eruption and mechanical energy of eruptions. Tuff rings evolve through a stage of explosion breccia emplacement to a stage dominated by base surges which deposit thinly bedded layers. Tuff cones may be built when continuing activity evolves into a third stage, 52characterized by rock emplaced by poorly inflated base surges and ballistic fallout. The differences are related to water: melt ratios; fragmentation of melt attains maximum explosive energy when the water:melt ratio is about 0.5 for basaltic compositions, whereas initial ("vent-coring") eruptions with small ratios result in the formation of breccia with abundant cognate and accidental fragments. Increasing ratios cause development of expanded dilute surges which deposit thin-bedded layers as tuff rings. Higher ratios produce "wetter" and denser eruption columns giving rise to poorly expanded surges, hence dominantly massive beds and tuff cones. The rates of magma and water influx controls the process, and such "cycles" may be interrupted, reversed or alternate. Most commonly, tuff cones may have craters filled or partly filled with lava, and agglutinated spatter and cinders. In some volcanic fields, scoria cones contain deposits of phreatomagmatic origin commonly developed during their initial eruptive stages. Littoral Cones Littoral Cones Littoral cones are mounds of hyaloclastic debris constructed by hydroclastic explosions at the point where lava enters the sea, and represent craters that lack feeding vents connected to subsurface magma supplies ("rootless") and form where lava or pyroclastic flows move over small ponds of water, swamps, springs or streams. The cones commonly occur as crescent-shaped ridges, breached by the source lava or more rarely as complete cones with craters occurring above lava tubes. Explosion centers are near or at the shore line, therefore about half of the radially exploded material falls into the sea, leaving a half-cone on land. A typical littoral cone is characterized by: (l ) a wide crater and low rims; (2) steep inner slopes exposing truncated strata unconformably mantled by in- dipping strata; and, (3) gentle outer slopes merging with the slope of the underlying terrain. Littoral cones are typically composed of hundreds of very poorly sorted, poorly defined beds ranging from a few centimeters to over 10 cm thick. They consist of fine- to coarse-grained ash, lapilli, and angular blocks up to l .5 m and bombs to l m in longest dimension. Ash > 4ø is commonly no more than 5% of the total ash content, and is composed of sideromelane, microcrystalline basalt and broken phenocrysts. Some layers contain accretionary lapilli and bedding sags, suggestive of abundant water vapor in the explosion clouds. Shield Volcanoes Shield Volcanoes Shields are mainly confined to volcanoes produced by rapid accumulation of fluid basaltic lavas. Three principal types of shields can be distinguished: the Icelandic, Hawaiian, and Galapagos types. Icelandic Shields These simple and most symmetrical shield volcanoes are formed entirely or almost entirely by effusive eruptions from central summit-vents. Icelandic shields range in height between 50 and 1,000 m, averaging 350 m. The angles of slope are unusually small, varying from 1° to 5°, but exceptionally may be as steep as 10°. The summit-craters or sinks of Icelandic shield volcanoes are approximately circular, and most measure less than l km across. They have raised rims built of spatter from lava fountains and by repeated overflow from lava lakes. Most summit-craters are essentially cylindrical pits with flat floors that may contain smaller collapse pits. A few collapse 53pits may also be found on the flanks of the shields, but they are distributed at random. There are few radial fissures and lines of parasitic cones on the flanks, and concentric fissures around the summit-vents are rare. Some flows traveled more than 25 km beyond the edges of the shields, even over slopes of less than l°. Virtually all lavas are of pahoehoe type and contain abundant small lava-tubes. Hawaiian Shields These shield consists of three superposed units. The lowest and largest is made up of pillow basalts and other products of submarine eruptions from deep-water vents. These are overlain by hyaloclastites formed by eruptions in shallow water and by subaerial lavas that entered the sea. The topmost unit consists of thin, subaerial flows. Profiles of Hawaiian shields during mature stages of growth tend to be slightly convex in their upper parts, uniform in their middle parts, and slightly concave in their lower parts. The upper slopes are only about 45 to 65 m per km, whereas the middle slopes are three to four times as steep (ca. 10° ), and around the base they diminish to 2° or 3°. During the mature, shield-building stages of growth, basalts are discharged at intervals of a few years from vents near the summit and along radial rift zones. Virtually all of the lavas are of the pahoehoe type near their source, but downslope many change to aa. As the shields grow, the flows tend to increase in average thickness and have more varied compositions. Pyroclastic ejecta constitute less than one percent of the volume, and most is produced in the last stages of development. Following the mature stage of growth, the volcanoes pass into a declining stage when the shields are buried beneath steeper-sided structures built of thicker, and more varied types of lava and pyroclastic debris, and by clusters of parasitic scoria cones and trachyte domes. During their initial stages of growth, Hawaiian shields may resemble Icelandic shields in that they are built by repeated outflows from central summit-vents. Subsequently, the summit activity is accompanied by an increasing number of eruptions from rift zones on the flanks. Mature Hawaiian shields are distinguished by prominent rift zones that converge at the summit-calderas. Main rift zones result from gravitational stresses in the subaerial structure and have no regional tectonic significance. Active eruptive fissures tend to begin within or near the calderas and then extend downslope, producing relatively short en echelon gashes. They commonly extend farther downslope than the lowest lava-vents, and the maximum outflow is not always from the lowest vents, but from the widest parts of the fissures. Locally, short en echelon eruptive fissures progress up- slope rather than downslope. Eruptive fissures may cross the floors of summit-calderas and collapse pits. Rift zones, as wide as 3 km, are marked at the surface by collapse pits, open cracks, small grabens, and chains of cinder- and spatter- cones, and are underlain by swarms of subparallel vertical or steeply inclined dikes. Galapagos Shields Shield profiles in the mature stage of growth resemble those of giant tortoises or overturned soup bowls with deeply indented tops. Visible parts measure between 15 and 30 km across at sea level, and they rise to heights of approximately 1,100 to 1,700 m. The slopes of some are less than 20°, but the middle slopes reach angles of 15° to 35°, flattening rapidly near the base and just as rapidly to wide benches surrounding the summit-calderas. Some fissures are gaping cracks that show little or no vertical displacement of the walls; others are marked by chains of small scoria- and spatter-cones. Lava flows pour from the fissures, and also issued from radial fissures on the flanks of the shields, where they are accompanied by chains of scoria-cones, generally breached on their lower sides. Galapagos shields may have started growth by discharge from a summit-vent; subsequently, mainly by eruptions from circumferential fissures surrounding summit-calderas and 54only partly by eruptions from radial fissures. Pyroclastic ejecta are present only during the declining stages of growth. Nordlie has presented a model for the origin of the Galapagos shields. Composite Cones Composite Cones Cones built partly of lava flows and partly of fragmental ejecta have long been referred to as stratovolcanoes. It is preferable to refer to these cones as composite. External Form The shapes of the cones are influenced by their manner of growth. During early stages, most eruptions issue from central conduits. During late stages, discharge of lava tends to take place more and more from radial fissures far down the flanks, while explosive eruptions may continue from the summit. Radial fissures result from the tumescence of cones and from the increasing hydrostatic pressure of magma in the lengthening, central pipes. Explosive eruptions from summit- craters, combined with effusive eruptions from fissures on the flanks, are partly responsible for the typical concave profiles. Coarse ejecta blown from summit-vents accumulate close to their source and their deposits generally have high angles of repose, whereas fine ejecta, which accumulate chiefly at lower levels, are distributed over a larger area, and so tend to flatten the lower slopes. Erosion is a major factor in accentuating the concave profiles, by removing debris from the upper slopes and depositing it around the base. In general, the older a cone and the longer it has been extinct, the more pronounced is the concavity of its slopes. The shapes of large composite cones are also influenced by: (a) the composition of the erupted magmas, (b) differences in the ratio of lavas to pyroclastic ejecta, (c) the depths of explosion-foci within the conduits, and (d) the location, size, number, shapes, and inclinations of eruptive vents. Composite cones composed almost wholly of lava flows tend to develop forms intermediate between those of domes and shield volcanoes, whereas others, composed almost wholly of pyroclastic ejecta, tend to have symmetrical forms with uniform slopes. Composite volcanoes have a conical form only if built mainly or exclusively by eruptions from a central, more or less cylindrical conduit. Those built by eruptions from elongate fissures or subparallel fissures have forms that resemble those of overturned canoes. Elongate forms may also result from progressive migration of the main conduit along a fissure system. The basal parts of volcanoes that rise from a shallow sea floor differ from those built wholly on land in that they contain much more fragmental material, both volcaniclastic debris and sediments, and are intruded by many flat-lying sills. Internal structure Dissection reveals essentially conformable, outward dipping layers of lava and pyroclastic debris. In multiple-vent cones, there may be many angular unconformities, owing to overlap of lavas and pyroclastic debris from different vents. Dikes, sills, and central plugs are commonly exposed by deep erosion. The dikes may be irregular in trend and thickness, but most are approximately radial. Swarms of subparallel dikes are exceptional; so are ring dikes and cone- sheets. Sills are commonly mistaken for lava flows. 55Fillings of the central conduits, usually referred to as plugs, necks, or pipes, vary greatly in size and shape, as well as in composition, structure, and texture. Few plugs exceed a few hundred meters in diameter, though some measure more than a kilometer across. They may pinch and swell, but most merge downward into stocks. Crudely cylindrical shapes are typical, but many plugs are markedly elongate or star-shaped. Typically they contain massive medium- to-coarse- grained rocks in differing stages of alteration. Growth Sequences Many large composite exhibit an evolutionary sequence, and consist of superposed or coalescing structures built by discharge of magmas of different composition and by eruptions of different kinds. Such volcanoes are compound and grouped genetically into normal, recurrent, and inverse types: 1. Normal: Cones in which successively younger, generally smaller forms are built by eruption of increasingly differentiated magmas. Many large “normal” composite cones are composed principally of basalt and andesite while younger products erupted on the flanks form domes of dacite and rhyolite or basaltic cinder cones and flows. 2. Recurrent: Cones display a repetition of eruptive sequences and magma types, and include "Somma-type volcanoes,' built chiefly of andesite, the tops of which collapse to form calderas as a result of large explosive eruptions of more siliceous magma. New composite cone then develop within the caldera and follow a similar evolutionary sequence. 3. Inverse: Cones in which the normal differentiation series from less to more siliceous magmas is reversed. Parasitic Cones The growth of parasitic cones on the flanks of large composite volcanoes is a sign of old age. Not uncommonly, these cones develop at successively lower levels as the volcanoes approach extinct. Usually, they are made up of more basic and more siliceous differentiates. Parasites may be concentrated along lines or belts that reflect structural trends in the subvolcanic basement, or in a crudely concentric arrangement. The concentric rings may reflect cone-sheets or ring dikes at depth. A crudely radial arrangement of parasitic cones and domes is much more common. The number of parasitic cones on most large composite cones is seldom more than ten or a dozen. 56IX. Craters, Calderas, and Grabens IX. Craters, Calderas, and Grabens These volcanic features, which represent negative landforms or depressions, are of several kinds and origins. Small to moderate sized depressions, when more or less circular in plan, are called craters. Larger quasi-circular depressions, which typically are many times greater than that of any included vents, are referred to as calderas. The lower size limit of calderas is a 1.5-kilometer diameter. Larger, elongate volcanic depressions include: (a) open fissures; (b) grabens formed by dropping of long narrow fault blocks, and, (c) larger down-faulted basins or troughs. The origin of these volcanic depressions can be conveniently divided into explosion and collapse related types. Explosion Craters Explosion Craters These are normally formed at the summit of lava, spatter, cinder, and ash cones, and of composite volcanoes. They result partly from the inability of the cone to build up over the vent, partly from collapse of the summit due to coring out of the cone by explosions or withdrawal of lava from the upper part of the conduit Collapse Craters Collapse Craters These are nearly circular craters that perforate the surface of a volcano without any surrounding debris cone. They result from sinking in of part of the volcano, and range from a few meters to kilometers in diameter. They may be few meters to over 300 meters deep, but are generally roughly circular in plan. The wall of the crater are nearly vertical during early stages of growth, but become more sloped with erosion. Calderas Calderas In form and origin, calderas resemble large pit craters, differing only in size. There floor is broad and generally flat. Curved faults typically lie outside the main caldera and are more or less parallel with the caldera walls. These walls are steep cliffs, formed by faulting, with banks of talus at the base formed by fragments falling from the cliffs. Outward-dipping lava flows in the caldera walls are typically cut off abruptly. There is generally no question of a collapse origin because only small amounts of pyroclastic material are present. The history of a caldera consists of a series of collapses and refillings. The entire floor may rise as a unit, with the entry of new magma from below. Hawaiian calderas, for example, appear to have grown by gradual coalescence of a number of pit craters that formed, one after another, on the summit of the shields. Fault scarps are more or less tangential to the group of pit craters, indicating sinking of the summit as a whole as well as local subsidence. During the late stages of activity, Hawaiian volcanoes tend to fill their calderas and build a cap of lava flows and pyroclastic cones at the top of the shield. The crater filling lavas can be easily recognized from the main shield lavas by: 57(a) thick beds that are essentially horizontal; and, (b) low vesicularity compared to the flank lavas. Calderas are also formed at the summits of composite volcanoes, some being up to 5-6 miles wide and more than 1000 meters deep. Formation of these calderas usually takes place late in volcanic activity, and often follows a long pause in activity during which the cone may be deeply eroded. Magmatic material ejected during the caldera-forming eruption is generally rhyolitic or dacitic in composition. Commonly ash flows appear to have been erupted from the same series of arcuate fissures on which the circular summit block subsides to form the caldera. Where formed on a composite volcano, the eruptive fissures are sometimes part way down the slope of the cone, but in other instances, the distribution of tephra indicates ejections came from restricted vents at the summit. The volume of material thrown out by explosions equals only a few percent of the caldera volume. It is clear that the caldera has not resulted from explosive decapitation of the mountain, but from subsidence along ring fractures. A huge volume of juvenile magma erupted indicates collapse probably occurs as partial drainage of an underlying magma chamber removes support from beneath the summit of the volcano. Such an interpretation implies the presence of a fairly large magma reservoir. Classification of Calderas Classification of Calderas Williams (1957) suggested that calderas fall into one of the following types: Krakatoan; Katmai; Valles; Hawaiian; Galapagos; Masaya; and Atitlán. Williams and McBirney further suggested that these types can be divided into groups distinct not only in the nature of the magmas with which they are associated, but also in geophysical properties: (a) Group I - moderate to strong negative gravity anomalies resulting from deep infill of light fragmental material (Krakatoan, Katmai, and Valles types); and, (b) Group II - moderate to strong positive gravity anomalies, resulting from the presence of very dense rock at shallow depths beneath the caldera (Hawaiian, Galapagos, Masaya, and Atitlán types). Krakatoan Type Krakatoan type calderas are formed by the foundering of the tops of large composite volcanoes following explosive eruptions of siliceous pumice from one or more vents or, in some instances, from arcuate fissures on the flanks. The volume of ejecta is usually much less than 100 km 3 . Prior to formation of the caldera, the site may have been occupied by a cluster of coalescing, composite cones. Eruption follows a long period of quiet, and may be immediately preceded by up to 16 years seismic activity. The climactic eruption probably lasts no more than 24 hours, but in that brief time, 2 to 6 km 3 of pumice are discharged. The eruptive column may rise more than 20 km, and ejection velocities may reach 500 m/sec. At first, most of it fall in showers, but some of the later deposits may be laid down by glowing avalanches. As activity progresses, explosions increase in vigor with a growing area of dispersion. Concluding eruptions may be influenced by inflow of groundwater that causes phreatomagmatic outbursts to form thin layers of vitric ash interbedded with accretionary lapilli, and simultaneous heavy rains convert some pumice deposits into lahars. Finally, lava may be issued at the foot of volcano. As a consequence of these eruptions, the top of the volcano collapses to produce a caldera. After collapse, intra-caldera domes of andesite and dacite rise at intervals from the floor 58The shape of the caldera may be been determined by three principal controls: 1) arcuate bays forming scalloped margins that may indicate the collapse of individual cones into cupolas above a common reservoir; 2) radial grabens extending outward from the caldera that may reflect extensions from the central reservoir, and, 3) structural trends of the pre-volcanic basement. Katmai Type Katmai type caldera collapse results from drainage of a central magma reservoir to feed new volcanoes or fissure eruptions beyond the base of the cone. The caldera may be up to 3 km across and up to 1200 m deep. This depression results partly from explosions, but mainly occurs when magma is drained from the central conduit to replace and mix with magma in a neighboring fissure system. Both magmas are discharged from fissures on the flanks. Valles Type Valles type caldera foundering takes place along arcuate fractures independent of preexisting volcanoes as a consequence of simultaneously with a discharge of colossal volumes of siliceous pumice, usually more than 100 km 3 . Calderas of this type, which include the world's largest, are produced by collapse accompanying and following discharge of great volumes of tephra from arcuate fissures formed by the rise of large subjacent bodies of siliceous magma. The caldera typically measures at least 20 by 25 km across, and its walls enclose a circular moat partly filled by a ring of rhyolite domes that may surround a resurgent dome (up to 650 m high and 10 by 13 km wide) produced by uplift of the caldera floor. The caldera does not occupy the beheaded top of an ancestral volcano, but rather, occupies the site of a group of cones and domes with related lavas and pyroclastic deposits ranging in composition from rhyolites to basalts. The total volume of ejecta is close to 200 km 3 . Foundering of caldera may temporarily close the arcuate eruptive vents. Pumice flows sweep outward from the arcuate boundary fissures, while others poured into the sinking caldera. Shortly after the caldera was formed, rhyolitic lavas and pumice may be discharged from fissures near its center and a lake may form on the floor. A group of late rhyolite domes are typically built over ring-fractures within the surrounding moat. Hawaiian Type Hawaiian type calderas are formed by collapse of the tops of shield volcanoes during late stages of growth. Prior tumescence is followed by subterranean drainage of basic magma from beneath the summit region into rift zones and, in many cases, by flank eruptions of lava. Calderas, up to 4.4 by 3.3 km across and up to 180 m deep, form at the summit when the shields has grown almost to its full height and while eruptions are still vigorous and frequent. As the calderas increase in width and depth, lavas pour from fissures cutting the floors and walls, and from rift zones on the flanks. Ultimately subsidence comes to an end, and the caldera-filling stage begins. Eruptions are spaced at much longer intervals and are more explosive. Flows and tephra accumulate inside the caldera until it is finally buried. The calderas typically have steep walls, interrupted in places by step faults, and are partly surrounded by benches up to 3 km wide traversed by inward-facing fault scarps. Both are also closely associated with pit craters; the 59calderas grow in part by coalescence of adjoining pits. It has been suggested that the calderas result from drainage of the central conduits into rift zones and by eruptions of lava far down the flanks of the shields. However, the volume of lava discharged during any eruption is generally much less than the volume of the summit-collapse and concurrent subsidence. The calderas typically collapse following broad tumescence of the shield when the rift zones are distended by magma, whether or not lavas are emitted at the surface. The fractures surrounding the calderas dip steeply inward. This, together with the basin-restricted distribution of the lavas inside calderas, suggest that the floor sinks as a wedge-shaped block, indicating that subsidence of the summit takes place after fractures have been widened by general inflation of the shield. Galapagos Type Galapagos type caldera are also formed by collapse during late stages of growth of basaltic shields but engulfment results from injection of magma and eruptions of lava from circumferential fissures near the summit and less frequently from radial fissures on the flanks of the shields. Galapagos shields are characterized by concentric fissures around the calderas and radial fissures on the flanks. The calderas are up to 7 by 10 km wide and up to 845 m deep. The number and length of the arcuate fissures vary and no single fissure describes a full circle at the surface. There may be at least four concentric fissures, the outermost more than a kilometer from the rim. There is little vertical displacement on any of these circumferential fissures, which probably owe their origin to periodic distentions of the shields by rising magma. Radial fissures on the flanks of the shields seem to have discharged fewer lavas than the rim fractures around the calderas. Masaya Type Masaya type calderas are formed by piecemeal cauldron subsidence of a broad shallow depression occupying most of the central portion of a low inconspicuous shield; eruptions from arcuate and radial fissures outside the caldera play no part, and nearly all the lavas are contained within the boundary scarps. The cauldron occupies most of the central part of a low basaltic shield, and measures up to 6 by 11 km across. Its outer slopes are gentle. No arcuate vents border the rim and no rift zones cut the flanks of the volcano. The walls, which range up to 150 m in height, consist of basaltic lavas and scoria. Craters and scoria-cones on the floor of the depression may arranged in a circle, and may discharged most of the lavas on the floor and probably outline an early collapse structure. Lava lakes occupy the active vents intermittently, their level fluctuate rapidly, and locally, their disappearance is followed by localized collapse. The crater floor may subside as much as 200 m during a single eruptive cycle. The scalloped margins of the depression suggest a succession of roughly cylindrical collapses resulting from periodic migration of magma underground. Explosive eruption and flank flows play no role in the development of the depression. Atitlán Type Atitlán type calderas are formed by cauldron subsidence unrelated to an earlier cone but associated with eruptions from volcanoes near the rim or from nearby fissures. The cauldron is caused by intermittent collapse, possibly resulting from underground migration of magma to feed adjacent volcanoes. Fault scarps that enclose one side of the basin may be deeply eroded and from 300 to 600 m high, whereas on the other side, may become fresher, steeper, smaller, and pass under the summit of a caldera-rim stratovolcano. 60Cauldrons Cauldrons The name cauldron is often applied to calderas formed by passive subsidence into a broad, shallow magmatic reservoir. The dimensions of a cauldron approaches or in some cases exceeds that of its associated cones, and many cauldrons form where there has been no large volcano. There is no sharp demarcation between calderas and cauldrons; surface eruptions and foundering into a shallow magma chamber contribute to the development of both. The principal difference is that calderas are associated with a withdrawal of magma, while cauldrons are thought to result from a passive foundering of the roof of a static or rising body of magma. Volcano-Tectonic Depressions Volcano-Tectonic Depressions Volcano-tectonic depressions are bounded by faults of tectonic origin and may reach dimensions of tens or even hundreds of kilometers in length. These great down-faulted troughs are often, though not always, formed on the crests of broad arches. For example, the Taupo Basin on the North Island of New Zealand is more than 60 km long and 30 km wide and contains several large composite volcanoes. The basin formed by collapse of the crest of a broad arch, forming a series of grabens, At about the same time, great masses of rhyolitic ash flows erupted from fissures in the crest of the arch or from faults and fractures bounding the grabens. The composite volcanoes only formed within the graben after collapse. Thus, these depressions are genetically similar to Krakatoan and Valles-type calderas. Resurgent Calderas Resurgent Calderas As already mentioned, volcanic activity commonly continues after caldera collapse. The caldera may eventually be completely filled and obliterated by later volcanic rocks. Renewal of activity in Valles-type calderas appears to be commonly associated by the up-bowing of the caldera floor, sometimes by thousands of meters. These calderas are referred to as resurgent calderas. Their updomed floor is stretched and cracked, with grabens commonly formed across the dome. Later eruptions are localized along the grabens and along the ring fractures that bound the caldera. 61X. CLASSIFICATION OF VOLCANIC X. CLASSIFICATION OF VOLCANIC ERUPTIONS ERUPTIONS Volcanic eruptions have been classified on several different bases. The usual ones are based on the form of the eruptive vent, the location of the eruptive vent, and the character of the eruption. Nature of Vent Nature of Vent The most generic classification is that based on the form and location of the vent, and involves: (1) fissure eruptions - volcanic material issues from a crack or fissure. These eruptions tend to be typical of most oceanic volcanoes. (2) central vent or pipe eruptions - volcanic material issues from a vent at the apex or center of the volcano. These eruptions, which tend to be typical of most continental volcanoes, are subdivided into several types: a) summit; b) flank; c) lateral; and, d) adventive. Eruptions from vents near or beyond the base of the mountains are sometimes called excentric. Those eruptions that occur within the summit crater are often called terminal eruptions, whereas flank eruptions near the summit are referred to as subterminal eruptions. Terminal and subterminal eruptions produce no marked lowering of the top of the magma conduit in the conduit, but lateral and excentric eruptions generally result in a marked lowering of the magma column and an increase in gas activity at the summit. The later eruptions, which build parasitic or adventive cones, tend to be wholly effusive in nature while explosive gas release occurs at the terminal vent. Styles of Eruptive Activity Styles of Eruptive Activity Volcanic eruptions have been classified on several different bases. The usual ones are based on the form of the eruptive vent, the location of the eruptive vent, and the character of the eruption. In the latter case, the eruptive style is based on the physical nature of the magma, the character of explosive activity, the nature of effusive activity, the nature of dominant ejecta, and the structures built around the vent. Based on these criteria, six principal eruptive styles, most named after the volcano which best exemplifies that type of activity, can be recognized. Hawaiian Eruptions These eruptions of basaltic, highly fluid lavas of low gas content give rise to effusive lava flows and less voluminous pyroclastic debris. Most eruptions start from fissures, commonly as a line of lava fountains, that ultimately coalesce to one or more central vents. Fragmental ejecta normally precede discharge of lava flows. Thin, fluid lava flows can gradually build up large broad shield volcanoes. A lava lake may be present in the summit crater of related shield volcanoes. Volcanoes with this activity include Mauna Loa and Kilauea-Iki, Hawaii. The characteristic features of these eruptions are: 62Physical nature of magma: fluid, basaltic. Character of explosive activity: Weak ejection of very fluid blebs; lava fountains. Nature of effusive activity: Thin, often extensive flows of fluid lava. Nature of dominant ejecta: Cow-dung bombs and patter; very little ash. Structures built around vent: Spatter cones and ramparts; very broad flat lava cones. Strombolian Eruptions These eruptions are characterized by discrete explosions separated by periods of less than a second to several hours in magma columns near the surface. Lava fountains are small and exceptional. Explosion clouds seldom rise more than 500 m, and are usually grey, with little or no lightning. Lava flows may follow explosive activity, and may continue uninterruptedly for months or years. Volcanoes with this activity include Stromboli and Mount Etna, Sicily, Italy; and Paricutin, Mexico. The characteristic features of these eruptions are: Physical nature of magma: moderately fluid. Character of explosive activity: Weak to violent ejection of pasty fluid blebs. Nature of effusive activity: Thick, not extensive flows of moderately fluid lava; flows may be absent. Nature of dominant ejecta: Spherical to fusiform bombs; cinder; small to large amounts of vitric (glassy) ash. Structures built around vent: Cinder cones Peléean Eruptions These eruptions, which are typically violent and destructive, involve glowing avalanches of fresh, effervescing magma. Separation of a gas cloud from the avalanche produces a nuée ardent that may move independently of the associated ash flow. Airfall ejecta are not widespread. Viscous magma follows to form steep-sided domes and spines or short, thick flows, the flanks of which may collapse by gravity or internal explosions to produce hot block-and-ash flows. Volcanoes with this activity include Mount Pelée, Martinique; Mount Mayon, Philippines; Santiaguito, Guatemala; and Mount Lamington, Papua New Guinea. The characteristic features of these eruptions are: Physical nature of magma: viscous; dacitic, andesitic, rhyolitic. Character of explosive activity: moderate to violent ejection of solid or very viscous hot fragments of new lava; commonly with glowing avalanches. Nature of effusive activity: domes and/or very short, thick flows; may be absent. Nature of dominant ejecta: Essential, glassy to lithic, blocks and ash; pumice. Structures built around vent: Ash and pumice cones; domes; local development of volcanic spines. Plinian Eruptions These eruptions, commonly lasting several hours to about 4 days, involve high eruption rate, voluminous, gas-rich eruptions that produce widely dispersed sheets of pyroclastic material derived from high eruption columns. The energy and characteristics of the eruption depends on: gas content of magma, rheology, vent radius and shape, and volume of magma erupted. Volcanoes with this activity include Vesuvius, Italy; Valles Caldera, New Mexico; and Mount St. Helens, Washington. The characteristic features of these eruptions are: 63Physical nature of magma: viscous; felsic (rhyolitic, trachytic, phonolitic, dacitic) becoming more mafic during course of eruption. Character of explosive activity: paroxysmal ejection of large volumes of ash, with accompanying caldera collapse. Nature of effusive activity: ash flows, small to very voluminous (up to 1,000 cubic km) sheets; may be absent. Nature of dominant ejecta: glassy ash and pumice; well-sorted; welding absent. Structures built around vent: Widespread pumice lapilli and ash beds; generally no cone building. Vulcanian Eruptions These eruptions are similar to Plinian, but are characterized by more explosive activity that produces a mushroom-shaped eruption cloud. Activity generally begins with phreatic (water-rich) eruptions that discharge lithic debris from the solid filling of the conduit. During the main phase, eruption of viscous, gas-rich magma forms a dark eruption cloud charged with vitric (glassy) ash. The tephra are well bedded, show marked gravity sorting, and generally, have widespread distribution. Volcanoes with this activity include Vulcano, Lipari Islands, Italy; and Okmak, Aleutian Islands. The characteristic features of these eruptions are: Physical nature of magma: viscous; basaltic to rhyolitic. Character of explosive activity: moderate to violent ejection of solid or very viscous hot fragments of new lava. Nature of effusive activity: Flows commonly absent; when present they are thick, and stubby; ash flows and base surges rare. Nature of dominant ejecta: Essential, glassy to lithic, blocks and ash; pumice; breadcrust bombs. Structures built around vent: Ash cones; block cones; block and ash cones. Surtseyan Eruptions These eruptions, also known as phreatomagmatic eruptions, represent violent explosions caused by rising basaltic (or more rarely andesitic) magma coming into contact with abundant, shallow groundwater or surface water. Broad pyroclastic cones of primarily ash, called tuff rings, are built by explosive disruption of rapidly cooled (quenched) magma. Magma fragmentation occurs chiefly due to thermal shock when the magma is quenched; particles are bound by fractures and broken vesicles. Volcanoes with this activity include Capelinhos; Surtsey, Iceland; and Taal, Philippines. The characteristic features of these eruptions are: Physical nature of magma: viscous; basaltic. Character of explosive activity: violent ejection of solid, warm fragments of new magma; continuous or rhythmic explosions; base surges. Nature of effusive activity: short, locally pillowed, lava flows; lavas may be rare. Nature of dominant ejecta: lithic, blocks and ash; often accretionary lapilli; spatter, fusiform bombs and lapilli absent. Structures built around vent: tuff rings Pyroclastic products of Surtseyan eruptions can be distinguished readily from those of Strombolian- or Hawaiian-type eruptions (into which they commonly grade as water is denied access to the magma during the course of the eruption). The criteria are as follows: 64Strombolian/Hawaiian Surtseyan Median diameter on or near cone 1/2 to > 16 mm usually 2 to < 1/8 mm Spatter common absent Particle shape achneliths common achneliths absent Accretionary lapilli absent common Impact structures (bomb sags) rare or absent common Thickness of individual beds > 1 cm, most > 5 cm commonly < 1 cm to 1 mm Alteration always oxidized; never oxidized; palagonite rare palagonite typical 65Appendices Appendices A. Pyroclastic Fall Deposits A. Pyroclastic Fall Deposits Pyroclastic fall deposits show: (a) more or less exponential decrease in thickness and grain size with distance from vent (b)mantle bedding (c) block impact structures (d)good to moderate sorting Exception - water-flushed ash may show (a) only but gives independent evidence for water flushing (e.g. accretionary lapilli, vesicles, water-splash microbedding). Fall deposits can be sufficiently hot to show primary welding when they accumulate near vent. 66B. Pyroclastic Flow Deposits B. Pyroclastic Flow Deposits Pyroclastic flow deposits show: (a) ponding in depressions, with a nearly level surface; (b)irregular thickness variations with distance from vent; (c) minimal sorting or internal stratification; (d) evidence for being hot (e.g. welding, pervasive thermal coloration, carbonization of contained plant remains, uniform direction of thermoremanent magnetization of contained clasts) Exception - low-aspect ratio ignimbrites include a mantling layer which passes laterally into the valley-pond ignimbrite. Note 1: ignimbrite can be defined as a pyroclastic flow deposit made mostly of pumiceous material (pumice, shards) Note 2: primary mudflows (lahars) resemble pyroclastic flow deposits but lack (d) 67C. Pyroclastic Flow Deposit Characteristics C. Pyroclastic Flow Deposit Characteristics Deposit Descriptions Ignimbrite Pumice and Ash Unsorted ash deposits containing variable amounts of pumice rounded salic lapilli and blocks up to 1 m in diameter. In flow units pumice fragments can be reversely graded while the lithic clasts can show normal grading; upgraded flows are as common. A fine-grained basal layer is found at the bottom of flow units. They locally contain fossil fumarole pipes and carbonized wood. The coarser smaller volume deposits usually form valley infills while the larger volume deposits may form large ignimbrite sheets. Locally, they may show one or more zones of welding. Scoria and Ash Topographically controlled, unsorted ash deposits containing basaltic to andesitic vesicular lapilli and scoriaceous ropy-surfaced clasts up to 1 m in diameter. They may in some circumstances contain large non- vesicular cognate lithic clasts. Fine-grained basal layers are found at the bottom of flow units. Fossil fumarole pipes and carbonized wood may also be present. The presence of levees, channels and steep flow fronts indicate a high yield strength during transport of the moving pyroclastic flow. Vesicular Andesite and Ash Topographically controlled, unsorted ash deposits containing intermediate vesicular (between pumice and non-vesicular juvenile clasts) andesite lapilli, blocks and bombs. Fine-grained basal layers, fossil fumarole pipes and carbonized wood may also be present. Block and Ash Topographically controlled, unsorted ash deposits containing large, generally non-vesicular, jointed, cognate lithic blocks which can exceed 5 m in diameter. The deposits are generally reversed graded. Fine-grained basal layers, fossil fumarole pipes and carbonized wood may be present. Surface manifestations include the presence of levees, steep flow fronts and the presence of large surface blocks, all of which indicate a high yield strength during transport of the moving pyroclastic flow. 68D. Pyroclastic Surge Deposits D. Pyroclastic Surge Deposits Pyroclastic surge deposits show: (a) draping of topography (b)rapid and irregular or periodic thickness fluctuations (c) general decrease in thickness and grain size with distance from source (d)commonly erosional base Two main types of pyroclastic surges occur: A - cold, damp or wet ... base surges; deposits show: (a) good internal stratification or cross stratification (b)great grain size variations between contiguous beds (c)evidence for dampness (e.g. accretionary lapilli, vesicles, plastering of up-vent side of obstacles) (d)association with vents in low-lying or aqueous situations or vents containing water (crater lakes) B - hot, dry ... ground surges and surges of nuée ardent type; deposits show: (a) little or no internal stratification (b)good sorting, depletion in fine or light-weight particles (c) evidence for being hot Exception - very similar deposits underlying ignimbrite can be produced by sedimentation from the pyroclastic flow. 69E. Pyroclastic Surge Deposit Characteristics E. Pyroclastic Surge Deposit Characteristics Deposit Description Base Surge Stratified and laminated deposits containing juvenile vesiculated fragments ranging from pumice to non-vesiculated cognate lithic clasts, ash and crystals with occasional accessory lithics (larger ballistic ones may show bomb sags near vent) and deposits produced in some phreatic eruptions which are composed totally of accessory lithics. Juvenile fragments are usually less than 10 cm in diameter due to the high fragmentation caused by water/magma interaction. Deposits show unidirectional bedforms. Generally, they are associated with maar volcanoes and tuff rings. When basaltic in composition they are usually altered to palagonite. Ground Surge Deposits are less than 1 m thick, composed of ash, juvenile vesiculated fragments, crystals and lithics in varying proportions depending on the constituents in the eruption column. Typically enriched in denser components (less well vesiculated juvenile fragments, crystals and lithics) compared to accompanying pyroclastic flow. They show unidirectional bedforms; carbonized Generally wood and small fumarole pipes may be present. Ash Cloud Surge Stratified deposits found at the top of and as lateral equivalents of flow units of pyroclastic flows. They show unidirectional bedforms, pinch and swell structures and may occur as discrete separated lenses. Grain size and proportions of all components depend on the parent pyroclastic flow. Can contain small fumarole pipes. 70