G E O . 4 1 6 V O L C A N O L O G YG E O . 4 1 6 V O L C A N O L O G Y

I. Physical Nature Of Magmas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4Structural State of Silicate Melts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5

Controls on Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5Silica composition . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6Temperature . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6Time . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 6Volatiles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7Pressure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7Crystal content . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 7Bubble Content . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8

Yield Strength . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8Specific Heat . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8Thermal Conductivity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8Density . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9Electrical Conductivity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9Seismic Wave Velocities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 9

II. Generation, Rise And Storage Of Magma . . . . . . . . . . . . . . . . . . . . . . . . . . 10Nature of Crust and Upper Mantle . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 10Heat Sources . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11

Mechanisms of Melting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11Partial Melting . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 11

Segregation and Rise of Magmas Through The Mantle . . . . . . . . . . . . . . . . . . . . . . . 12Rise of Magmas Through Brittle Lithosphere . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13Flow of Magma . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14

Flow Rates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 14Nature of Flow Regime . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15Flow Instabilities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 15

High-Level Reservoirs and Subvolcanic Stocks . . . . . . . . . . . . . . . . . . . . . . . . . . . . 16

III. Eruptive Mechanisms . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18Opening Of Vents . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18Mechanisms of Explosive Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 18Nature of the Gaseous Eruptive Column . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 19Bubble Nucleation And Growth . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20

Pressure Relations . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 20Ejection Of Pyroclastic Material . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21

Ejection Velocities . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 21Eruption Energy . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 22

V. Lava Flows . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23Volume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23Length and Thickness . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23

Velocity of Flow . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24Discharge Rates . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 24Physical Properties of Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25

Temperature and Cooling of Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 25

Morphology Of Lava Flows . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26Pahoehoe Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26

External Structures of Pahoehoe Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26Internal Structures of Pahoehoe Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 27

Aa Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28External Structures of Aa Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 28

Block Lava . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29Internal Structures of Blocky Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 29

Pillow Lava . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 30

VI. Volcanic Domes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31External Features of Volcanic Domes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 31Internal Structures of Volcanic Domes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 32

VII. Products Of Volcanic Explosions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33Terminology and Classification . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33

Origin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33Fragment Size . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33

Airfall Ash Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 35Dispersal . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36Structures Of Airfall Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 36Morphology of Ash Particles . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37

Pyroclastic Flow And Surge Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 37Relationship to Topography . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 38Flow Units and Cooling Units . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 39Components . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40

Characteristics of Ash-Flow Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40Internal Layering . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 40Gas-Escape Structures . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41Textural Relationships . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 41Segregation of Crystals and Lithics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42Temperature Effects . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42Welding and Compaction . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 42Structures Related to Temperature and Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . 43

Classification and Nomenclature of Pyroclastic Flows . . . . . . . . . . . . . . . . . . . . . . . 43

VII. Laharic Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47General Features . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47

Surface of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47Basal Contact of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48Components of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48Grain-Size Distribution . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48Grading . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 48Fabric . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49

Origin of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 49

VIII. Structures Built Around Volcanic Vents . . . . . . . . . . . . . . . . . . . . . . . . . 50Cinder Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50

External Form . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50

Internal Structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51Maar Volcanoes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 51Littoral Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52Shield Volcanoes . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52

Icelandic Shields . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 52Hawaiian Shields . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53Galapagos Shields . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53

Composite Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54External Form . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54Internal structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54Growth Sequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55Parasitic (adventive) Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55

IX. Craters, Calderas, and Grabens . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56Explosion Craters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56Collapse Craters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56Classification of Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57

Krakatoan Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57Katmai Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58Valles Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58Hawaiian Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58Galapagos Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59Masaya Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59Atitlán Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59

Cauldrons . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60Volcano-Tectonic Depressions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60Resurgent Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 60

X. Classification Of Volcanic Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61Nature of Vent . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61Styles of Eruptive Activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61

Hawaiian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61Strombolian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62Peléean Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62Plinian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62Vulcanian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63Surtseyan Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 63

Appendices . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65A. Pyroclastic Fall Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 65B. Pyroclastic Flow Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 66C. Pyroclastic Flow Deposit Characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 67D. Pyroclastic Surge Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 68E. Pyroclastic Surge Deposit Characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69

G E O . 4 1 6 V O L C A N O L O G YG E O . 4 1 6 V O L C A N O L O G Y

I . P h y s i c a l N a t u r e O f M a g m a sI . P h y s i c a l N a t u r e O f M a g m a s

Magma is a completely or partially molten natural substance, which on cooling, solidifiesas a crystalline or glassy igneous rock. It is usually rich in silica and capable of flowing undermoderate differential stress. Magmas may carry rock fragments or crystals in suspension, and theynormally contain gaseous (volatile) components in solution.

Volcanic magmas fall within a strictly limited compositional range that reflects the physicaland chemical processes responsible for their generation and differentiation. Our concern is thephysical phenomena of volcanism, interpretation of which requires some knowledge of physicalproperties of magmas.

Unfortunately, we have only a meager knowledge of liquid properties. Much of what isknown can be explained in terms of the properties of Silicon (Si) and Oxygen (O) ions, which areusually the most abundant components. Si has a high charge (+4), small ionic radius (0.39 Å), andlow coordination number with oxygen (4 oxygens surround each silicon, forming the corners of atetrahedron). This results in strong ionic field strength and bonding with oxygen compared to othercations: Ca, Mg, Fe, Mn, Ti, Na or K. Al, which has similar but not as strong properties, plays asimilar role to Si in both liquids and crystalline solids.

Structura l S ta te o f S i l i ca te Mel t sStructura l S ta te o f S i l i ca te Mel t s

Modern concepts of silicate liquid structure are based on the Zachariasen Model. The atomsare bonded by forces similar to those between atoms of crystals, but lack long range periodicityand 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; theSi and Al cations are termed "framework cations." Other cations enter the melt in limited amountsas independent ions occupying positions between tetrahedra, and modify the basic structuralframework 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 20cation percent before the basic framework breaks down into smaller geometric units. In breakingliquid continuity into smaller units, the framework changes from an extensive network oftetrahedra, 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 separatetetrahedra not directly linked to each other.

Melt structure controls the physical properties of a magma. Viscosity is the most importantof these properties, because it plays a role in factors controlling both the style of volcanic eruptionand the physical nature of volcanic products.

5

V i s c o s i t yV i s c o s i t y

Viscosity is a fluid's internal resistance to flow. It represents the ratio of shearstress to rate of shear strain applied to a layer of thickness Z and permanentlydeformed in a direction x parallel to the stress. Mathematically, viscosity is expressedby:

s = s o + ηdmdt

n,

where s is the total shear stress applied parallel to the direction of deformation; so is the yieldstrength of the fluid or the stress required to initiate flow; η is the viscosity, expressed in unitscalled poises (dyne sec/cm2); dm/dt is the gradient of velocity dx/dt or strain rate over a distance Znormal to the direction of shear; and, n is an exponent which has a value of 1.0 or less dependingon 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> so) 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 < so) is applied, the substance is deformed elastically andreturns to its original form after the stress is removed. Some fluids do not require application ofsome initial force before they are permanently deformed by shear stress parallel to the direction ofshear. Such fluids are said to exhibit Newtonian behavior when n equals 1.0 and so equals zero.

Highly polymerized or non-Newtonian fluids (known as Bingham liquids) have a finiteyield 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 theycontain approximately 20% crystals. Liquids with suspended solid particles may have a non-linearrelation 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 viscousbehavior of complex silicate liquids, such as magmas, is difficult to predict, because nocomprehensive theory explains the effects of major cations or temperatures of magmaticconditions.

It is possible to estimate the viscosity of a magmatic liquid at temperatures well aboveliquidus temperatures (that is, temperatures at which only liquid is present) from chemicalcompositions and empirical extrapolation of experimental data on the linear relationship between hand temperature in simple chemical systems. The range of temperatures of naturally flowingmagmas, however, is near or within the crystallization interval, where stress-strain relationships

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are not linear (that is, they are crystal-liquid mixtures and show Bingham behavior). Under suchconditions, the only way to predict viscosities is by analogy with similar compositions investigatedexperimentally.

Silica composition

The strong dependence of viscosity of molten silicates on Si content can be illustrated bythose of various Na-Si-O compounds:

Na:Si:O η (poises)0:1:2 1010

1:1:2.5 282:1:3 1.54:1:4 0.2

The decrease in viscosity can be attributed to a reduction in the proportion of framework silicatetrahedral, 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 whichmagmas flow. Low temperatures characterize silica-rich rhyolite magmas, whereas the highesttemperatures are observed in basalts. Magmas do not crystallize instantaneously, but over aninterval of temperature. Few magmas, however, have a wide enough range of crystallization toremain mobile at temperatures far below those at which they begin to crystallize or much hotterthan those temperatures.

Temperature has a strong influence on viscosity: as temperature increases, viscositydecreases, an effect particularly evident in the behavior of lava flows. As lavas flow away fromtheir source or vent, they lose heat by radiation and conduction, so that their viscosity steadilyincreases. 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 adistance 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 manyhours before it reaches a steady value. The viscosity increases with time results partly an increasingproportion of crystals (which raise the effective magma viscosity by their interference in meltflow), and partly from increasing ordering and polymerizing (linking) of the framework tetrahedra.

7

Volatiles

The solubility of gases in magmas varies with pressure, temperature and composition ofboth the gas and the magmatic liquid. Because the volume of a melt with dissolved gas is less thanthat of a melt and separate gas (vapor) phase, solubility increases as gas pressure increases. Atconstant gas pressure less than total pressure, any increased load pressure on the melt lowerssolubility, 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 componentgenerally decreases with temperature, except possibly at high pressure. Consequently it is difficultto predict how volatile content of magma varies with depth. Nevertheless, it has been shown that atconstant temperature, solubilities of water in magmas with different compositions are notsignificantly different.

Nearly all magmas can contain more water or gases at depth than they can continue to holdin solution when they reach the surface. Basalts, however, normally contain less water thanrhyolites 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, H2S, SO2, CO,and CO2) on magma viscosity. No doubt, the effect of dissolved water is to lower viscosity, theeffect being greater for silica-rich than silica-poor magmas:

Magma T (°C) ηdry (poises) ηwet (poises)Rhyolite (~70% SiO2) 785 1012 106 (5% H2O)Andesite (~58% SiO2) 1000 104 103.5 (4% H2O)Basalt (~48% SiO2) 1250 102 102 (4% H2O)

Dissolved water disrupts the framework of linked Si and Al tetrahedra, but where suchpolymerization is already minor or absent, there is little effect. F and Cl are though to considerablyreduce magma viscosities; in contrast, CO2 increases polymerization, and therefore viscosity, inmelts by forming CO3-2 complexes.

Pressure

The effect of pressure is relatively unknown, but viscosity appears to decrease withincreasing pressure at least at temperatures above the liquidus. As pressure increases at constanttemperature, the rate at which viscosity decreases is less in basaltic magma than that in andesiticmagma. The viscosity decrease may be related to a change in the coordination number of Al from 4to 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:

η = ηo(1 - RC)-2.5

8

where η is the effective viscosity of a magmatic liquid, C is the volume fraction of suspended

solids; ηo is the viscosity of the magmatic liquid alone; and, R is a constant with a best-estimatedvalue of 1.67.

Bubble Content

The effect of gas bubbles (vesicles) on the bulk viscosity of magmas can be variable, anddepends 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; inbasaltic magmas, the bubbles may enhance the already low temperature and composition controlledviscosity. Rhyolitic magmas have high viscosities irrespective of the degree of vesiculation, andonly effect of high bubble content will be to reduce mechanical strength of the melt.

Yie ld S trengthYie ld S trength

Most magmas have an appreciable yield strength, which shows a marked increase belowtheir liquidus temperature. As yield strength increases, the stress required to initiate and sustainflow becomes greater, and the magma's apparent or effective viscosity is also increased.

Spec i f i c HeatSpec i f i c Heat

The specific heat (Cp) of magma, which is the heat required to change the temperature ofthe liquid 1 degree Celsius, is typically about 0.3 cal. gm-1. The specific heat contrasts greatly withheat of fusion or crystallization, which is the heat that must be added to melt or removed tocrystallize a unit mass that is already at a temperature where liquid and solid coexist. Heats offusion are typically about 65-100 cal. gm-1 at 1 atmosphere. Consequently, about the same amountof heat is involved in crossing the crystallization interval, as in raising or lowering the temperatureof the rock or liquid through 300°.

Thermal Conduct iv i tyThermal Conduct iv i ty

Igneous rocks and liquids are poor conductors of heat. Thermal conductivity depends ontwo 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 structureexpands. For rocks, the two effects balance each other up to their melting range. At hightemperatures, the thermal conductivity of mafic rocks normally declines at an increasing rate up to1200°C, above which, radiative heat transfer increases as does total thermal conductivity. Moresilica-rich rocks show increasing thermal conductivity at lower temperatures.

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Dens i tyDens i ty

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 ofhigher concentrations of Fe, Mg and Ca cations in basalts. In contrast, magma density decreaseswith increasing temperature and gas content. These densities increase a few percent between liquidand crystalline states.

The temperature dependence of magma density is given by the coefficient of thermalexpansion, about 2-3 x 10-5 deg-1 for all compositions. The pressure dependence of magmadensity is given the compressibility or fractional volume change, ∆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 cm2

dyne-1.

Elec tr i ca l Conduct iv i tyElec tr i ca l Conduct iv i ty

Electrical conductivity, which is low in pure silica melts, increases with increasingabundance of metallic cations, especially alkali elements, and increases abruptly in the meltingrange.

S e i s m i c W a v e V e l o c i t i e sS e i s m i c W a v e V e l o c i t i e s

Compressional or P-wave velocities are about 6 km sec-1 up to the melting range, thendecrease abruptly to 2.5 km sec-1 at higher temperatures. Shear or S-wave velocities are about 2-3km sec-1, which drop abruptly at melting temperatures.

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I I . G e n e r a t i o n , R i s e A n d S t o r a g e O f M a g m aI I . G e n e r a t i o n , R i s e A n d S t o r a g e O f M a g m a

The subsurface processes by which magmas are generated and rise toward the surface areextremely complex. Before examining these processes, it is worthwhile to review what is knownconcerning the Earth's interior.

Nature o f Crust and Upper Mant leNature o f Crust and Upper Mant le

Most of what is known concerning the Earth's interior comes from geophysicalmeasurements, 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 severaldepths interpreted to represent discontinuities in the composition or structural state of minerals. Themost 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 ρ and the elastic properties (bulk modulus

K and rigidity or shear modulus µ) by the following expressions:

Vp = { [K + (4/3)µ]ρ}

Vs = (µ/ρ)1/2

The elastic properties are poorly known, but making certain assumptions, it appears that densityincreases to about 3.4 gm/cc at depths around 70 km, remains constant between 3.45 and 3.63 tothe base of the Low Velocity Zone. Both pressure and temperature increase with depth. Thetemperature 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 measuredin the crust are projected downward, they rapidly approach temperatures for beginning of meltingin 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 mustdiminish with depth.

Heat SourcesHeat Sources

Existence of magma indicates that at some depth beneath the Earth's surface, temperaturesmust be high enough to induce melting. One major problem associated with understanding thegeneration of magmas is the source of heat necessary to cause melt production. It is believed that

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the 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 lowabundances in probable mantle rocks, too low to yield through their radioactive decay the heatnecessary to generate magmas. Moreover, it can be shown that the melting process scavenges theseelements, 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 inducemelting within the Earth:

(a) Stress Relief: Pressure on the source region is released during tensional orcompressional 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 conditionswhere lowest temperatures on the solidus coincide with phase changeboundaries.

(c) Convective Rise : The source material rises by solid-state convection into apressure-temperature regime appropriate for melting

(d) Perturbation: A local decrease in thermal conductivity or density leads toheating or diapiric rise of the source material.

(e) Mechanical Energy Conversion To Heat: Force required to move one rocksurface over another without grinding and deformation converted to heat,because of thrust faulting, subduction, a propagating crack or flaw in theEarth's lithosphere, shear or Tidal energy dissipated in the solid earth.

(f) Compositional Change: The addition or subtraction of material changes therock composition to a new composition whose solidus lies at a temperatureless than the ambient temperature.

Partial Melting

Rocks are a heterogeneous assemblage of minerals, and each mineral is characterized by aunique melting temperature. Melting commences at grain boundaries, usually where three crystalsof minerals with the lowest melting temperatures meet. As melting progresses, channelwaysdevelop between grains. Temperatures probably never are high enough to completely melt thesource rock, and only part of or some of the minerals melt. This process is therefore called partialmelting.

Because of mechanical constraints, it is generally believed that at least 1-5% melting isrequired for the melt to separate from the unmelted (refractory) solid (crystalline) material. Meltingprobably never exceeds 35% because of the gravitational instability of low density liquid withhigher density refractory minerals. The composition of a partial melt (magma) depends on themelting conditions present in the Earth:

(a) temperature; (b) pressure;

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(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/orbrittle (crust) domains within the Earth. The manner in which magma rises differs between thesetwo domains.

Segregat ion and Rise o f Magmas Through The Mant leSegregat ion and Rise o f Magmas Through The Mant le

Several mechanisms of magma rise through the mantle have been visualized. Theseprocesses include:

(a) Deep Segregation: The melt forms along a dendritic network of joints andfractures in the zone of melting, and feeds into a smaller number of layertributaries eventually forming a larger channel at higher levels. With meltingconcentrated along grain boundaries, melt migration is caused by a thermal orpressure gradient or by capillary effects. This migration the presence of acritical proportion of melt before solid/liquid separation occurs. Two factorswhich 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 temperaturemuch higher than its surroundings, as it is cooled by adiabatic expansion andconduction to the wall rocks. If the liquid rises slowly through rocks that arebelow their melting temperature, the magma would crystallize quickly. Thus,magmas can only ascend once the temperature of their wall rocks have beenelevated, and successive batches of magma must tend to follow paths ofearlier 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 localizedbulges under the heavier layer. The low density layer moves upward at anaccelerated 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 densitydifferences, and the viscosity of the overlying rocks. Little or no separationof melt occurs in the zone of melting. Instead, the crystal-liquid mush risesand separation occurs at shallow levels. There again must be a delicatethermal balance between the diapir and its surroundings. Otherwise, itcrystallizes.

(c) Zone Melting: A body of magma rises by melting its roof, while it crystallizeson its floor. The zone of melting rises without actual movement of liquid andwith little loss of heat. Heat used in melting is regenerated by release of latentheat of crystallization. It has been estimated that a body of magma 7 km thickstarting at a depth could rise to within 8 km of the surface before crystallizingin about 1 million years.

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Rise o f Magmas Through Br i t t l e L i thosphereRise o f Magmas Through Br i t t l e L i thosphere

It is difficult to determine the level at which the lithosphere deforms by brittle fracture ratherthan by plastic flow - a depth represented by earthquake foci. There is strong evidence, in the formof individual and swarms of dikes, that large bodies of magma are tapped within the crust at a levelwhere rocks can fail by dilational fracture. However, temperatures and pressures in the vicinity oflarge magma bodies are not normally consistent with purely brittle fracture. The manner in whichmagmas rise through the lithosphere may be:

(a) Dilational Rise: This proposed mechanism by which magma may riseinvolves: (i) entrance of melt in fractures, and rise due to gravitationalbuoyancy; (ii) The fracture becomes extended vertically and/or horizontallyalong a plane normal to the minimum stress; and, (iii) The fracture closesbehind the magma as it passes and pressure on the wall falls below theconfining pressure, rebounding due to viscoelastic deformation. Such amechanism may explain the limited duration of basaltic fissure eruptions andthe apparent arrival of discrete batches. Many instances, however, exist whereacid or volatile magmas have apparently risen as pipe-like intrusions with littleor no evidence of horizontal deformation.

The ability of a magma to rise through brittle lithosphere is usuallyexplained in terms of depth and density contrast with the overlying rocks. Ifthe 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 depthof 50 km, the lithostatic pressure can exceed the pressure of a vertical magmacolumn enough to segregate liquid and cause it to rise. If the heights to whichmagmas can rise is solely dependent on the depth to source and a densityequilibrium, it would be expected that magmas with deep sources woulderupt at higher elevations, and vice versa. This is obviously not the case asdemonstrated by volcanoes of the Mexican volcanic belt.

More important limitations to magma rise are probably the heatcontent, and rates of ascent and cooling, which in turn, depend on the size ofthe magma body. Another important factor is the stress regime, whichgoverns the form of the intrusive bodies. The three basic magma stressregimes 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 andhorizontal stresses are low and about equal on the surrounding rocks, themagma conduits tend to be cylindrical. Thus, the form taken by a magmabody may change drastically during its ascent. It is likely that near thesurface, a cylindrical pipe is the most efficient form of conduit, because flowvelocity increases and heat losses decrease as the horizontal section increases

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in size and becomes equidimensional. Thus, conduits tend to becomecentralized at the intersection of two or more fracture systems.

(b) Non-Dilational Rise: As mentioned previously, there is ample evidence thatsome magmas have forcibly displaced rocks into which they have intruded,but others have made room for themselves by stoping or elevating the roofrocks. It is obvious that the critical elements are heat, and the manner inwhich the magma crystallizes, the shape and size of the body, and the volatilecontent of the magma.

An excellent example of non-dilational rise is illustrated by theformation 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 mechanismsmay be capable of boring through the Earth's crust and creating diatremes:

(i) Highly energized gases of deep-seated origin bore through thecrust, opening channelways for the rapid ascent of magma; or,

(ii)Explosive eruption is triggered by vaporization of heatedgroundwater propagated downward as pressure is released onprogressively deeper gas-charged horizons.

F l o w o f M a g m aF l o w o f M a g m a

Knowing the rheological or fluid properties of magmas, we might be able to apply basicfluid dynamic principal to predict flow regimes of intrusive and extrusive magmas under variousphysical conditions. Unfortunately, a rigorous approach to our understanding of flowcharacteristics is not currently possible in the face of incomplete information about essentialparameters of specific cases. Nevertheless, some insight into magma ascent processes may begained by considering simple examples and approximations.

Flow Rates

The volumetric flow rate of a viscous fluid through a cylindrical channel under a constantpressure gradient is given by:

Q = (ΡΠr4)/8ηL

where Q is the volume flow rate in cm3 sec-1, P is the pressure drop in bars, r is the channel radiusin cm, η is the viscosity of the fluid in poises, and L is the length of the channel in cm. Applyingthis relationship to a large (about 200 km3) simple funnel-shaped magma chamber which is filledwith basaltic magma (η = 300 poises) via a 3-km-long, 200-m-wide, cylindrical feeder pipe at itsbase and a pressure drop through the pipe of 1000 bars (1 kb/3.3 km), we find:

Q = (3.14 x 1000 x 1016)/(8 x 300 x 3 x 105) = 4.36 x 1010 cm3/sec or 3.76 km3/day.

15

This simple calculation is important in that it illustrates that movement of large quantities of magmain 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 ofinterest. For example, in the case of an initially heterogeneous magma, the liquid would becomeeffectively homogenized by turbulence. The conditions that determine laminar or turbulent flow canbe determined by calculating the dimensionless Reynolds number, Re, which in terms of averageflow rate is given by:

Re = (2ρQ)/ηr Π

where ρ is the density of the fluid. Turbulent flow occurs when Re > 2000. For the previous

example, with ρ = 2.6 gm/cm3,

Re = (2 x 2.6 x 4.36 x 1010)/(3.14 x 104 x 300) = 2.39 x 104

Hence, flow of the basaltic magma within the conduit would be turbulent. The higherviscosity of acid magmas, however, renders turbulent flow unlikely in these cases. Because theviscosity of magmas normally exceeds 103 poises and velocities are rarely greater than a fewcm/sec, flow is probably laminar under most geologic conditions.

It can be expected that the non-Newtonian characteristics of magma also have an effect onflow behavior. Because a certain yield strength must be exceeded before many magmas can bedeformed by viscous flow, velocity gradients in the margins of a moving magma are likelydifferent from those of more familiar liquids like water.

Shear stress in the boundary of the moving liquid is greatest near a stationary surface anddiminishes toward the interior. Thus, if viscosity is uniform throughout the entire flow width, thenthe velocity distribution is parabolic. But if heat is lost at the stationary boundary and the effectiveviscosity increases sharply with falling temperature, the flow profile is more arcuate. Thesedifferent flow profiles reflect both the effect of falling temperature on both viscosity and yieldstrength of the magma.

In many cases, it is likely that a zone of static liquid will form a layer between the movingliquid and its solid boundary. Heat transferred from a cooling magma to surrounding wall rocksalso 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 theliquid large enough to produce gravitational instability, the liquid overturns and free convectionaccelerates the rate of heat transfer. The onset of convection in an infinite horizontal layer ofviscous fluid having an upper and lower surface is given by the dimensionless ratio of buoyant toviscous forces known as the Rayleigh number, Ra:

Ra = (L4 αTgβ)/ηK

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where L is the height of the layer in cm, αT is the coefficient of thermal expansion, g is the

constant of gravitational acceleration (980 cm/sec), β is the vertical temperature gradient in K cm-1,

η is the kinematic viscosity (η⁄ρ), K is the thermal conductivity of the magma in cal gm-1 K -1, and

ρ is the fluid density. Ra for a vertical tube heated from below is given by the same expression,except that L4 is substituted by r4 where r is the characteristic radius of the tube in cm.

The critical Ra value above which convection begins is about 1700, approximately the samevalue calculated for magmatic bodies of most common shapes. For a magma body of given sizeand viscosity, the principal variable is thermal gradient, β, a function of heat loss to the top orsides of the magma body. For Ra < 1000, transfer of heat is predominantly by conduction; steadyconvective heat transfer sets in at approximately Ra > 10000, and strong eddying motion is attainedwhen Ra = 100000. Bodies with thickness or radius greater than 10 m are likely to convect if theirheat 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, butquite small bodies having high heat flux values, should also be quite unstable.

High-Leve l Reservo ir s and Subvo lcan ic S tocksHigh-Leve l Reservo ir s and Subvo lcan ic S tocks

The erosion of extinct volcanoes reveals the presence of simple and multiple stocks ofmedium- to coarse-grained rocks. Generally, the stocks are 1- to 10-km-wide, circular to oval incross-section, and grade upward into a maze of inward dipping sills, steep radial dikes, and conesheets. Most of these intrusive rocks have made room for themselves by stoping rather thanforcible 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 andequilibrate in such shallow reservoirs prior to eruption. Other than what we see within deeplyeroded volcanoes, however, little is known concerning the volcanic reservoirs beneath activevolcanoes, except what is indicated by geophysics:

(1) Seismic methods: These methods have been used to detect large magmabodies at depth because of the inability of the Shear or S seismic waves to betransmitted through liquids. The distribution of earthquakes generated withinor 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 surroundsan aseismic zone, which may represent one or more bodies of magma beneathKilauea. Several types of earthquakes of volcanic origin are recognizedaccording to the location of their foci and the nature of earthquake motion:

(a) A-type volcanic earthquakes: These earthquakes take place in andbeneath volcanoes at places deeper than 1 km, generally in the rangefrom 1 km to 20 km. They are generally less than 6 in magnitude.

(b) B-type volcanic earthquakes: These earthquakes originate usually inand adjacent to active craters at extremely shallow depths. The

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magnitudes are generally extremely small. The earthquake motionsconsist mainly of vibrations with periods in the range of 0.2 sec. to1.0 sec.

(c) Explosion earthquakes: The maximum amplitude or magnitude of theearthquake has a close relationship with the intensity of explosiveeruption and is approximately proportional to the kinetic energy ofthe eruption. The earthquake motions show a predominance oflonger wave length as compared with those of the A-type volcanicand tectonic quakes. The associated detonations or air vibrations ofexplosive eruptions are remarkably strong.

(d) Volcanic tremors: Earthquakes take place incessantly or continuouslywith a short interval, such as every several seconds, so that motionsare recorded continuously. These earthquakes may originate fromextremely shallow positions in or near the crater, or at deep levels(20-30 km at Kilauea). Various wave forms are found in volcanictremors, including surface waves of Rayleigh and Love type.

(2) Gravity Measurements: Precise gravity measurements may also reveal thepresence of an anomalous mass of magma at depth, and provide a means ofconstructing subsurface structural models. Gravity surveys have shown thatthe Hawaiian volcanoes have crudely cylindrical cores composed of denserock only a few km below their summits. Gravity measurements have alsosuggested the presence of large batholith-size, low-density bodies of magmaor intrusive rock beneath many large calderas. They also indicate that Cascadevolcanoes lie within grabens, or down-dropped tectonic blocks, underlain bysimilar subvolcanic intrusions.

(3) Infrared Radiometry: This technique is used to detect the presence of bodiesof rock or magma at elevated temperatures.

(4) Tiltmeter Measurements: Precise leveling and tilt measurements have beenused to detect deformation caused by the intrusion of magma into shallowlevels. Such measurements have been used to estimate the depth andgeometry of the intrusions, because they provide precise informationconcerning the horizontal as well as the vertical components of movement.

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I I I . E r u p t i v e M e c h a n i s m sI I I . E r u p t i v e M e c h a n i s m s

Opening Of VentsOpening Of Vents

Rare observations indicate that during the initial phases of a volcanic eruption: (i) thefractures through which magma reaches the surface represent planes of dilation propagated aheadof slowly rising magma; (ii) the appearance of lava is preceded by a mild release of steam or heatedgroundwater; 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 thedepth of magma origin. Few explosive events are singular in nature, but rather represent an erraticsuccession of surges.

Magma does not reach the surface unless it is sufficiently heated to remain fluid and topenetrate the overlying barrier of cold rocks and groundwater. In order for these conditions to bemet, it appears that a minimum conduit width and flow rate of magma within the feeder dikes isrequired. The final ascent of magma to the surface is neither sudden nor violent, but rather is asteady 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 thelithosphere. Eruptions cease not because of a lack of magma, but due to a reduction in pressure.

Mechani sms o f Exp los ive Erupt ionsMechan i sms o f Exp los ive Erupt ions

All explosive eruptions involve the sudden release of energy by gas under pressure, but theway gas expansion acts on magmas varies widely. The explosivity of a volcanic eruption does notcorrelate directly with either volatile or silica content of the magma alone: the lowest is in those ofolivine basalts, but highest in those of basanites and lamprophyres. The major factors whichdetermine 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 theyvesiculate. 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 increasesas 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 diffusingthrough a narrow boundary layer, through convective processes or by vesiculation and rise of

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bubbles. Once a magma becomes saturated, it may rise and reach a level at which the pressuresexerted by the overlying rocks are low enough to permit vesiculation.

Expansion accelerates the rise of magma, so that the pressure of the overlying rock columnis reduced at a faster rate, and eruption ensues. This process by which a reduction of lithostaticpressure 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 byinjection of hotter magma (remember that, in both cases, a resulting temperature increase decreasesgas solubility).

In most cases, the initial phases of eruption result in the ejection of gases and disruptedmagma or ejecta with in a gas-charged cloud or eruption column.

Nature o f the Gaseous Erupt ive Co lumnNature o f the Gaseous Erupt ive Co lumn

To understand fully eruption mechanisms it is useful to examine the characteristics of theeruption column and how it varies as magma reaches to the vent:

(a) Temperature Relations: Exsolution and expansion of gas significantly coolsmagma as it rises. If there is good thermal equilibration between the magmaand gas, the extent of cooling can be very great, e.g. there can be 300°Ccooling of a vesiculating basaltic magma, if it expands adiabatically from thepressure at which gas exsolution begins. The temperature of the gas is largelydependent on the proportion of the two phases, and the efficiency of the heatexchange. The latter is strongly dependent on size because only ejecta ormagma fragments less than 5 mm can attain thermal equilibrium with the gasduring an eruption; silicate particles therefore account for most of the heat. Ifthe source of the gas is meteoric water, the heat used to flash the water tosteam tends to buffer the temperature eruption at around 100°C. As theeruption column emerges from the vent, it continues to cool as it expands andmixes with air.

(b) Density Relations: The density of the eruptive column influences its capacityto carry fragments suspended in the gas stream. The smaller particles aresubject to drag forces larger than their inertial forces, and thus, have lowerterminal velocities so that they behave like gas particles. Particles less than0.1 mm in diameter have so low terminal velocities compared to the velocityof the gas stream, that they contribute to the effective density and viscosity ofthe eruption column. A greater proportion of fine particles therefore enhancesthe ability of the eruption column to support large clasts or fragments.

(c) Viscosity Relations: A marked increase in magma viscosity occurs as a resultof falling temperature and reduced water content during eruption. As aconsequence, there is a slower expansion rate of bubbles as the magmaapproaches the surface. Conversely, the increased proportion of gas lowersthe overall viscosity if the gas phase becomes large enough to be continuous.

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Bubble Nuc lea t ion And GrowthBubble Nuc lea t ion And Growth

In order to understanding the mechanisms of explosive eruptions, it is useful to considerthe manner in which gas exsolves from the magmatic liquid. Even in the most viscous magmas, therate of bubble nucleation is very high. In order to evolve and grow, gas bubbles must reach aninitial size that balances the surface tension (σ) of the magma at the gas-liquid interface.

The pressure of gas inside the bubble acts over a cross-section Πr2, and is balanced by

surface tension around the circumference of its walls in the same cross-section (2Πrσ). Therefore,

the gas pressure must exceed a value of 2σ/r before it can expand. Stable micron-sized bubbles canform if the gas pressure is greater than 6 bars (dry) or less (water-saturated).

Phenocrysts (large suspended crystals) accelerate vesiculation because bubbles that nucleateon 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. Theexsolution of water vapor increases surface tension to different degrees in different magmas, whichmay explain why bubbles tend to expand intact in some magmas but coalesce in others. Exsolutionand 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 ofresidual gaseous phase, when the magma approaches the surface. There are four components ofpressure in the vesiculating magma:

(a) the pressure of the overlying magma column: (ρgh)

(b) the pressure required to drive the magma through the conduit:P = 12Vηh/r3 in cylindrical conduits

P = 12Vηh/w in fissure conduits

where V is the magma flow velocity, h is the length of the conduit, and r isthe conduit radius or w is the fissure width.

(c) the pressure required to overcome surface tension: The essential condition isthe relationship between gas pressure in bubbles to the strength of thesurrounding liquid. The strength of a vesiculating magma may be determinedby the bubble density: when the proportion is low, it is an important factor,but as the proportion increases, surface tension becomes important. Theforce of surface tension acting around the circumference of each bubble exertsa pressure over the cross-sectional area of the bubble, so that the totalpressure from surface tension through the vesiculated liquid is:

P = 2n2/3σ/r

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where n is the number of bubbles per unit volume and r is their averageradius. The excess pressure of the gas phase, ∆P, exerts a force per unit ofcross-sectional area of vesiculating magma, and must be greater than:

∆P > 2n2/3σ/r + τ

where τ is the critical tensile stress of the magma. For porosities greater than50 percent, this excess pressure need only be a bars in order forfragmentation of the magma to occur.

(d) The pressure required for the bubble to expand against the viscous resistanceof the surrounding liquid:

P = 4η/ r (dr/dt),

where (dr/dt) is the expansion rate of the bubbles. This pressure, whichvaries between 10-2 bars and several hundred bars, is strongly dependent onmagma viscosity. In a fluid basaltic magma, a bubble with a 1 cm radius cangrow radially at a rate of 0.5 mm/sec, more than enough to accommodate gasexpansion at low pressure, but in viscous magmas, the expansion rate is twoto three orders of magnitude slower and the pressure buildup is greater. Thefinal sizes and gas pressures of bubbles are mainly a function of magmaviscosity: the effect of increased viscosity during exsolution arrests expansionwhen 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 lattercomponents are small. After the magma has vesiculated to the point that it behaves as acompressible fluid, i.e. the gas forms a continuous phase in which silicate liquid is carried insuspension, the second component, the dynamic pressure, becomes dominant.

Ejec t ion Of Pyroc las t i c Mater ia lEjec t ion Of Pyroc las t i c Mater ia l

As mentioned previously, the ability of the eruption column to carry in suspension andeject fragments of disrupted magma is determined by the column density. The nature of ejecta andthe 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 arepart 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 thegas stream. The ejection velocity is the difference between the velocity of the gas stream and thevelocity with which fragments would settle under static conditions. The minimum ejection velocity

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can be estimated from the maximum distance that blocks of a given size travelled from the vent toimpact:

R = V2sin2J/g

where R is the distance, V is the initial velocity and J is the ejection angle. R has a maximum valuewhen J = 45°. The ejection angle is seldom as low as 45°; ejection angles tend to be 80° or moreabove the horizontal, and increase with depth to the focus of explosions. Velocities calculated withthis expression are less than the actual ejection velocity, because as soon as a block leaves theeffect of the gas stream, air resistance reduces its range, especially when it is small and has littlemomentum.

For a given velocity, moreover, the ejection distance varies directly with the mass of theblock, and inversely with its drag coefficient and cross-sectional area. The drag coefficient varieswith the shape, surface roughness, and velocity of block, and with the viscosity and density of theatmosphere. For a given initial velocity, large blocks travel farther than smaller ones, because theirinertia is higher, and momentum is less retarded by air resistance.

Below a few centimeters diameter, fragment movement is strongly retarded by wind andthermal currents. Estimated ejection velocities are on the order of 500-600 m/sec. Lower velocitiesare produced by the convective rise of warm air and gas. These currents, which are capable ofcarrying only fine dust, may reach great heights above the volcano, but velocities rarely exceed afew 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 fineparticulate material.

Eruption Energy

The energy release during a volcanic eruption is a summation of varied and often offsettingforces:

(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.

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I V . L a v a F l o w sI V . L a v a F l o w s

Lava flows are the products of extrusion of a coherent magma body onto the Earth'ssurface. The external forms and internal structures of lava flows are the result of both the physicalproperties of the magma and the external environment in which extrusion takes place. The principalphysical property that determines the nature of the lava flow is the magma viscosity, which is itselfinfluenced by both the chemical composition of the magma and its temperature. The rate of magmasupply to the flow is also important.

The external environment includes the steepness of the slope on which the lava isdeposited, and the presence or absence of water and/or ice.

VolumeVolume

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 themagma viscosity increases with increasing silica and alkali content. The volumes of most historicallava flows are generally measured in the 10ths or 100ths of cubic kilometers. Some of the largestknown lava flows include:

(a) 1669 Mt. Etna lava - ≈1 km3,

(b) prehistoric McCarty Flow (New Mexico) - ≈7 km3, and,

(c) 1783 Laki basalt flow (Iceland) - ≈12.2 km3.

All of these lavas are basaltic; siliceous rarely exceed 1 km3, with individuals some times only afew square feet in area and a few inches thick being known.

Length and ThicknessLength and Thickness

Because siliceous magmas are usually more viscous than basic ones, siliceous lavas tend tobe the shortest and thickest of all flows. Some lava flows are formed by a single gush of liquidspreading as a single unit. More frequently, it is found that repeated gushes of liquid have givenrise to intertonguing layers known as flow units. Subaqueous flow tend to remain fluid longer thanterrestrial flows because with increasing depth of water, exsolution of volatiles is suppressed andviscosity remains high due to dissolved water.

(a) Basaltic Lavas: Fluid basaltic lava flows in Hawaii extend for more than 35km with an average thickness of 5 meters. Some Icelandic basalts can betraced 80 km, whereas several Columbia River plateau basalts extended formor than 100 km from their source vents. One Columbia River basalt flowhas been traced over an area of 130 X 240 km, and has a thickness between30 and 50 meters.

The length of lava flows is determined largely by the magma effusionrate. A high effusion rate, where lava spreads rapidly from the vent, usuallyresults in a single flow unit. A low effusion rate, in contrast, results in lavasof limited extent that pile up layer on layer. It appears that basaltic lava flows

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that originate from fissures spread for distances that are roughly proportionalto the third power of their thicknesses.

(b) Andesitic Lavas: Andesitic flows generally have thicknesses of up to 30meters, and are usually 5 to 15 km in length. Pyroxene andesite lavastypically are more extensive than those of hornblende andesite. Hornblendeandesite 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 theselavas 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.

V e l o c i t y o f F l o wV e l o c i t y o f F l o w

The flow velocity of lava flows depends on a number of different factors: (i) rate ofeffusion, (ii) magma viscosity, (iii) volume of magma extruded, (iv) magma density, and, (v) theslope and nature of the channel in which it flows. As expected, flow velocity diminishes withdistance from the source. A pronounced velocity gradient exists within lava flows, extending fromthe middle toward the top, bottom and sides. Without a surface crust, the fastest movement occursin the upper and middle parts of the flow, but once a crust forms, the fastest-moving part movesincreasing downward into the lava. Some typical flow velocities are:

(a) Basaltic Lavas30-60 km/hr Hawaii8-75 km/hr Vesuvius

(b) Siliceous Lavasusually on the order of 10's or 100's of meters/hour.

Discharge RatesDischarge Rates

The discharge rate of lava flows from the volcanic vent depends principally upon thefluidity of the magma, and size and dimensions of the conduit. Like flow velocity, discharge ratesdecrease during the course of an eruption. Some typical basalt discharge rates are:

(a) 1947 Hekla - 75000 to 1250 m3/sec,(b) 1887 Mauna Loa - 5 million m3/hr, and,(c) 1946 Parícutin - 2 to 6 m3/sec.

The discharge rates of intermediate and siliceous lavas are generally much lower than those ofbasalt, but there are notable exceptions:

(a) Sakurajima - 1666 m3/sec, and,(b) Santorini - 45000 m3/sec.

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P h y s i c a l P r o p e r t i e s o f L a v a sP h y s i c a l P r o p e r t i e s o f L a v a s

Temperature and Cooling of Lavas

Most lavas are erupted at temperatures below their beginning of crystallization, and onlyrhyolitic obsidians are aphyric, or free of crystals. Because of their low thermal conductivity andhigh specific heat, most lavas are well insulated and cool slowly. Relatively little cooling takesplace through most of the course of the flow, especially if the eruption temperature is greater than1100°C.

The principal heat loss of a lava is through radiation from its surface. This can be expressedby the Stefan-Boltzmann equation:

Q = sT4,

where Q is the energy radiated per cm2/sec, T is degrees Kelvin, and s is the Boltzmann constant(5.67 X 10-5 ergs/sec/cm2/deg4). Because of the 4th power temperature relation, a small amount ofcooling greatly reduces the radiative heat loss. Only a minimal amount of heat may be conducted tothe air or ground, as indicated by:

Q = 2K(Ts-To)(t/(Πα))0.5,

where Q is the heat flux per unit time t, K is the conductivity of the ground, α is the thermalconductivity, Ts is the surface temperature, and To is the initial temperature of the ground. Owingto the low thermal conductivity and thermal diffusivity of soils and rocks, heat losses due toconduction are only a degree or two per hour.

Lava temperatures can be measured with (i) an optical pyrometer in which the color ofincandescent 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 colorof 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 maybe estimated from the relation:

η = gρsinAd2/3V

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where V is the mean velocity, g is the acceleration of gravity, A is the slope angle, d is the depth ofthe flow, and ρ 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 relationare low, because the velocities measured at the surface are greater than the mean velocity of theflow.

It is also possible to estimate lava viscosities from surface wavelengths of ripples in thelava crust using:

η = 2.61ρλ1.5

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 ofthe lava accounts for flow fronts and levees ceasing to flow laterally even though slope angles maybe great enough.

M o r p h o l o g y O f L a v a F l o w sM o r p h o l o g y O f L a v a F l o w s

Lava flows exhibit a variety of morphologies that depend on the magma viscosity and theexternal environment. Several types of lava flows are recognized to occur in lavas of different bulkcomposition.

Pahoehoe Lavas

These lavas are characterized by smooth, billowy, ropy or entrail-like crusts of quenchedlava. 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 upto 30 cm thick.

(c) Shelly: This very frothy lava has a minutely spinose sharkskin-like surface. Locally, aropy or corded surface develops when the fluid magma moves beneath a thin, partlycongealed crust, causing it to wrinkle and fold either convex downstream or inparallelism with the flow direction.

(d) Slabby: These pahoehoe lavas are characterized by broken crusts, forming slabs afew 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-scalesurficial 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 5-30cm in height. The coils develop along shear zones between relativelystationary and adjacent blocks, being moved by undercurrents.

(b) Lava Blister: A mound of continuous lava crust, a few mm to a m in heightand width, caused by the accumulation of gas beneath the lava surface.

(c) Tumulus: This dome-shaped structure, resembling an ancient burial moundand typically having an oval ground plan, forms as a result of upwelling of

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magma beneath a fairly thick lava crust. A tumulus forms when the lavabelow the crust is obstructed downstream. The tumulus may be up to 50 min length, and is locally 6- to 10-m-high. The tumulus surface is similar tothat of the surrounding flow, except that it is generally cracked radially. Lavaoften rises in the cracks to form either small unrooted lava flows, or bulbousmounds, up to a few m in height and width, called Squeeze-Ups.

(d) Pressure Ridges: These transverse, convex downstream, ridges representlava crust which has been heaved up into elongate mounds as much as 0.8km long and 50 m high. These ridges form as a result of the flow crust beingpushed against some obstacle by continued movement from behind. Elevationof the crust into an anticline is aided by the hydrostatic pressure of liquidbeneath the crust. In some cases, continued movement of the lava results inthe overturning of the fold with the gentle slope on the lava source side and asteep slope away from it. Locally, the folded crust breaks and slides forwardover the steep side of the ridge, forming a thrust fault. The crust of pressureridges is generally broken, and many of the ridges consist of a heap ofvariably oriented blocks.

(e) Hornitos, driblet- and spatter cones: These are small mounds or chimney-likespires that are built over eruptive vents or more commonly cracks in the lavacrust (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 aroundstanding trees. After flow level subsides, a hollow cylindrical column is leftby the carbonation of plant material.

Internal Structures of Pahoehoe Lavas

In addition to these external features, pahoehoe lavas may exhibit a number of internalstructures which include:

(a) Flow units: formed by the intertonguing of lava streams derived from thesame flow.

(b) Columnar Jointing: Contraction, the result of thermal stresses within thecooling lava, produces fractures that are propagated in a plane normal to thedirection of cooling. These fractures bound 5- or 6-sided, polygonalcolumns that develop perpendicular to the cooling surface. The columns,which vary from 5 cm to >3 m in width, are typically straight and haveparallel sides, but some may be curved. Throughout individual flow units,columns may variable considerably in dimension and cross-section, but athree-fold subdivision is typically recognized:

(a) upper colonnade(b) entablature(c) lower colonnade.

Invariably, the columns are cut by cross-joints, some curved upward ordownward as ball and socket joints. Discontinuous cooling leads to thedevelopment on the sides of the columns of chisel marks which mark theposition of isotherms during cooling. Although columnar joints are commonin all types of lava flows, it should be noted that they also characterize some

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pyroclastic ash-flows that have been emplaced at temperatures high enoughfor fragments to become welded together. They are also conspicuous in somesubvolcanic dikes (where they occur normal to the dike walls, stacked likefirewood when exposed by erosion) and in intrusive necks like Devil'sTower.

(c) Lava Tubes: These structures, which range from a few cm to 30 m or more indiameter and from a few km to 20 km in length, develop by the flow ofmolten magma within a confined interior channelway. In the upper parts of alava flow, migration of magma eventually becomes restricted to thesechannelways. Fast-moving flows are characterized by relatively straight lavatubes, whereas slow-moving flows tend to contain meandering and branchingtubes. If lava drains out of the tube before complete solidification, it leavesstrandlines on the tube walls. In cross-section, lava tube walls are marked byconcentric layers of congealed lava. Completely filled tubes show concentricbands of vesicles, platy joints parallel to the walls, and/or radiating jointcolumns. Lava stalactites and stalagmites may form by dripping of still-fluidlava from the tube ceiling; some may consist of sulfate minerals or opal.

(d) Pipe vesicles and spiracles: These gas cavities are formed when lava passesacross wet ground, generating steam. The steam bubbles rise into the lavaand form lines of vesicles or small tubes, usually less than 0.5 inches indiameter. If the upper end of the gas tubes are bent in the direction of lavamovement, they are called pipe vesicles, and have been cited as possibleindicators of flow direction. Where the steam bursts upward into the lava, itexplosively creates an irregular, up to 10 m diameter, cylindrical openingcalled a spiracle. The spiracle generally terminates within the flow rather thanextending through it, and may contain mud blown up from the underlyingground.

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. Basedon external form, various subtypes of aa lava can be distinguished:

(a) Aa Rubble flow: The lava crust consists of small, loose and semi-detachedfragments.

(b) Aa Clinker flow: The lava crust consists of loose and semi-detachedfragments that measure more than several cm in diameter.

(c) Furrowed aa flow: The lava is intermediate between aa and pahoehoe, with avery rough ropy surface that is locally arborescent.

Aa lavas flow like a caterpillar tread, dumping talus over the snout and then overriding their owndebris. 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 surficialstructures that include:

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(a) Lava Gutters: These channels develop when faster-moving parts of the flowdrain away from slower-moving parts and flow bottom as the supply of lavadiminishes or stops.

(b) Lava Levees: These longitudinal ridges develop by accretion of lava on theslower-moving parts or flanks of the flow, generally bounding the centralgutter.

(c) Lava Lobes: These features represent lava tongues that have generallydeveloped along flow margins after the levee is breached.

(d) Accretionary Lava Balls: These structures form, like snowballs, by the rollingup of solid fragments, either clinker or chunks, derived from the walls of theflow 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 thefragments have more regular forms and smoother faces. The surface blocks often approach cubesin form. Blocky lava flows form from more viscous lavas than aa flows, with the angular blocksformed by breaking up of the partly to wholly congealed upper part of the flow as still-mobilemagma moves beneath the thick crust. These flows are typically thicker that aa lavas (8-35 mthick), and fragmental material, which may constitute the entire thickness, makes up a greaterproportion of the flow than aa. The surfaces of blocky lavas are generally very irregular, withmany 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 tocomposition, may allow them to be distinguished from aa lavas:

(a) Ramp Structures: The high magma viscosity results in a large amount ofinternal shearing. Movement along the ground is retarded by friction,whereas moving liquid higher up in the flow tends to separate into a series ofsheets that slip over each other like a deck of cards. Movement of the lavasheets is predominantly parallel to the underlying surface. When solidified,these sheets may be very thin (few cm), and are defined by platy joints. Nearthe flow front, the extremely high viscosity of the magma may cause theshear planes to bend sharply upward. Ramps may be formed when localmovement upthrusts portions of the flow along the shear planes

(b) Lamination: These structures are formed by the upward bending of flowplanes and shear planes, often distinguished by different degrees ofcrystallinity. These laminations may form antiforms and synforms, the limbsof which may become crumpled.

(c) Spines: These structures form when a massive central part of the flow isprojected up into the fragmental portion of the flow, or even extended aboutit.

(d) Auto-Breccia: These deposits represent brecciation of the flow resulting fromshattering of the very viscous lava due to stress related to flowage.

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Pillow Lava

These flows are subaqueously extruded lava marked by bulbous forms. Pillow lavas mayform by the discharge of lavas into rivers, lakes, ponds or under glaciers, as well into oceans. Thepillow structures result from the protrusion of elongate lava lobes, which detach from and falldown the moving flow front. Lava pillows are often confused with pahoehoe toes, but the formerhave several distinguishing characteristics:

(a) Pillows are rimmed by chilled glass selvedges, formed by rapid cooling oflava by the surrounding water.

(b) Pillows vary from a few cm to several m in diameter, and are generallyspheroidal, ellipsoidal, or may be flattened in cross-section.

(c) Pillow tops are usually convex upward, whereas their undersurfaces may beflat, concave upward or project downward between the underlying pillows.

(d) Where gas cavities are present, these structures tend to be located within theupper part of the pillows.

Pillow lavas are locally found in association with several other types of subaqueous volcanicdeposits:

(a) Hyaloclastites (Aquagene Tuffs): These deposits are made up of brokenpieces of glass, formed by brecciation as a result of drastic chilling of thefluid lava.

(b) Pillow Breccia (Aquagene Breccia): These deposits are similar tohyaloclastites, but are dominantly composed of pillows and pillow fragments.

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V . V o l c a n i c D o m e sV . V o l c a n i c D o m e s

Over or near the vent, extremely viscous lava tends to pile up into steep-sided heaps ofmolten rock, known as domes or tholoids. Some domes also result from the bodily upheaval ofmaterial filling the upper part of the conduit. This semi-solid to solid material is pushed up like acork 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 oflava either internally or externally. Those domes that form by addition of lava through some formof extrusion through an opening in their crust, generally at the crest, are called exogenous. Morecommonly, lava squeezed up through the vent distends the mass above, so that these domes arecalled endogenous.

Most domes are composed of rhyolite, dacite or trachyte magma. Andesite domes are lesscommon, and basaltic domes are extremely rare. The size of domes varies greatly. Some are only afew meters across and a meter high, whereas the Mount Lassen dome is over 1 km across at itsbase and over 600 m high. Most domes are broader than high. In plan, domes are more or lesscircular, or very short ovals. Rarely are they elongate except as a result of extrusion of lavathrough a fissure vent. Where extrusion occurs as concentrations along a linear fissure, a row ofclearly separate and independent domes, sometimes with overlapping bases, may growsimultaneously. More commonly, domes grow successively along a fissure.

Most domes are short-lived features, because they are commonly destroyed by collapsepartly due to volcanic explosions and due to strains set during cooling. The speed of dome growthvaries considerably, but some rise by as much as 25 m/day.

E x t e r n a l F e a t u r e s o f V o l c a n i c D o m e sE x t e r n a l F e a t u r e s o f V o l c a n i c D o m e s

The exteriors of volcanic domes are distinguished by several distinctive geomorphicfeatures that develop at different stage of growth:

(a) Subsidiary Flows : The cooling outer part commonly ruptured by internal stretching, andlava oozes through the ruptures, forming trickles that move for varying distances downthe side of the dome. The degree to which a growing dome spreads out from the marginof the vent depends on the magma viscosity. Some domes spread very little, whereasothers 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 growingdome, and on the edges of the dome, break free and roll down the sides of the dome tocome to rest against its base. These accumulation of fallen crustal blocks form banks ofloose 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. Thedeposits are commonly massive, but sometimes show a suggestion of bedding due to thecrude parallel orientation of elongate fragments. There may also be occasional layer ofash, which accumulates on the surface of the growing dome during explosive activity.

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(c) Spines: These structures may form by squeezing out of viscous magma from withinruptures in a solid to semi-solid shell. The largest spine observed historically was thespine that formed during the 1902 eruption of Mt. Pelée, Martinique. This spine, whichhad a 9 month growth period, reached a height of about 300 m above the top of thesummit dome, and was 1000 m in diameter at its base. Without the loss of material due tocrumbling, 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 arethrust. The great spline of Pelée had a nearly vertical straight face and an opposite facethat was characterized by a curved surface that was polished and striated up the upwardmovement of the spine through the aperture in the solid carapace of the dome. Spines arecommonly accompanied in their development by huge, steep-sided pyramidal or conicalpeaks called pitons.

(d) Coronet Explosions: Volcanic explosions commonly occur around the base of domes andspines. The surfaces of separation between domes and the surrounding rocks, orbetween the spine and dome crust, constitute zones of weakness that allow gases toescape from below and within the dome. Jets of gas may issue from around the base ofthe dome like spikes around a crown, and thus are called coronet explosions. Explosionsat the base of the Chaos Crags dome, Mount Lassen at 300-600 years ago, resulted incollapse of part of the dome, and led to the formation of great avalanches that rushedabout 5 km down valley leaving behind a blocky deposit now known as the ChaosJumbles.

In terna l S truc tures o f Vo lcan ic DomesInterna l S truc tures o f Vo lcan ic Domes

Some endogenous domes have a series of concentric, onion-like layers, that result from thegradual expansion from the mass of a somewhat inhomogeneous magma. More commonly, domesare either essentially structureless, or show a divergent, fan-like structure in which ribs of the fanradiate upward from the vent. The fan-like structure is generally shown by aligned phenocrysts andlayers 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 issometimes expressed as concentric fractures on the dome surface. The internal structures of domeschange from horizontal to vertical upward from the vent.

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V I . P r o d u c t s O f V o l c a n i c E x p l o s i o n sV I . P r o d u c t s O f V o l c a n i c E x p l o s i o n s

Termino logy and Clas s i f i ca t ionTermino logy and Clas s i f i ca t ion

Fragments which are thrown out by volcanic explosions are referred to collectively asejecta, and accumulations of these fragments are known as pyroclastic rocks or tephra dependingon whether they are consolidated or not. Different sorts of volcanic explosions produce somewhatdifferent types of ejecta. Instantaneous or long-continued, weak or violent, all volcanic explosionsare the result of escape of gas from the magma, but they may be subdivided according to the originof 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 severaldifferent 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 moltenmagma or from rock that was already solid (non-magmatic ejecta). Non-magmatic ejecta mayrepresent:

(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 thesame eruption, are termed essential ejecta, and are typically partly or entirely glass (vitric). Type IInon-magmatic ejecta or fragments of older rocks formed during previous eruptions are termedaccessory ejecta, and most are partly or wholly crystalline (lithic). Some accessory ejecta consist ofcoarse-grained clots of several minerals that represent cognate material that was torn from theconduit walls or from parts of magma crystallized at depth.

Fragment Size

The most important classification of pyroclastic rocks is based on fragment size, althoughfragments 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:

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(1) Bombs, typically composed of low-viscosity basaltic magma, are ejecta largerthan 64 mm in average diameter that are thrown out of the vent in a moltenstate. Highly to moderately fluid, magma may be ejected both as long,irregular strips or as discrete blebs. Because they are fluid, their shape istypically modified during flight through the air and such fragments aretypically termed fusiform::

(a) Strips that break up into short segments form cylindrical or ribbonbombs, which are more less circular or flat in cross-section, and typicallyshow twisted, longitudinal fluting.

(b) Large blebs pulled up into spheres by the surface tension of the magmaform spherical bombs.

(c) Fragments that spin during flight form spindle- or almond-shapedbombs, characterized by longitudinal fluting, and one side smoother andbroader than the other. The smooth or "stoss" side represents the frontside as the bomb fell through the air, whereas the "lee" side is producedby frictional resistance of the magma dragging the still-plastic skin of thebomb toward this side. this resistance often forms a thin projecting rimalong the edge of the stoss side:

(d) Bombs of very fluid magma, that is projected only to moderate heightsand strikes the ground while still liquid, flatten or even splash to formpancake or cow-dung bombs.

At the other end of the spectrum, very viscous bombs are not rounded during flight, and althoughtheir outside is nearly solid, the inside is still plastic enough to expand as gases escape and producea skin that is broken by deep cracks, forming what is called bread-crust bombs. Most bombs aresimply an irregular and generally extremely vesicular lumps, which are described as cinder orscoria. Fusiform bombs may only form at the very end of an eruption, because they representdenser material than scoria or cinder, and they form at a stage when the amount of gas in themagma started to decrease.

Most bombs in cross-section are at least somewhat vesicular, and they are oftencharacterized by concentric layers of greater and lesser vesicularity. Exceedingly vesicular cinderare called pumice. Pumice of rhyolitic magma are characterized by vesicles that are stretched outinto long very thin tubes, giving the fragment a silky appearance. In contrast, far less abundantbasaltic pumice typically consists only of thin glass threads that mark the intersections of vesicles.These form the lightest rock (0.3 gm/cm3) known to exist, what is called thread-lace scoria orreticulate. 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 flattenand mod themselves to the underlying surface, forming an accumulation of flattened and weldedfragments called spatter or agglutinate. Masses of tephra containing a large proportion of bombs iscalled agglomerate.

Most bombs are less than 25 cm in diameter, but some may be exceptionally large, e.g. 6 mirregular, elongate bombs at Paricutin; up to 1 m fusiform bombs at Mauna Loa; and, up to 1.3 mcow-dung bombs at Stromboli (1965).

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(2) Blocks: These are angular ejecta larger than 64 mm in average diameter thatare thrown out of the vent in a solid state. Blocks typically are formed by thedisruption of the crust of a lava pool or a dome. They may be entirely cold orstill warm and incandescent when deposited. Accumulations of blocks arecalled breccia, and it is often desirable to specify whether the deposits arepyroclastic breccia or phreatic breccia depending on the type of volcanicexplosions responsible.

(3) Lapilli: These are ejecta between 2 and 64 mm in average diameter, may beessential, accessory or accidental in origin, and may be ejected in either aliquid or solid state. Lapilli are the most abundant type of fragment in cinderdeposits. Special forms characterize drops of basaltic lava that are ejected in avery 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 moistmaterial to a nucleus, like the growth of hailstones.

(4) Ash: This is tephra less than 2 mm in average diameter, may also beessential, accessory or accidental in origin, and ejected in either a liquid orsolid state. Depending on the material that composes the ash, it can bedescribed 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-sizedmaterial are referred to as ash layers, ash beds or ash blankets. Consolidatedash deposits are described as tuffs, or more specifically, depending on thepredominant constituent as:

(a) lithic tuff;(b) vitric tuff; or(c) crystal tuff.

Some ash deposits contain moderately to very abundant lapilli, blocks orbombs, and are termed lapillistone, lapilli tuff, tuff breccia or tuffagglomerate.

A i r f a l l A s h D e p o s i t sA i r f a l l A s h D e p o s i t s

Tephra produced by a volcanic eruption may be distributed by fall through the atmosphereor by flow over the ground surface. Tephra may also be dispersed by ocean currents, where ashhas fallen on seawater and coagulated. The most widespread pyroclastic product is airfall ashdeposits. Tephra of basaltic eruptions are much less voluminous than those of intermediate torhyolitic eruptions due to the less explosive style of basaltic volcanic activity.

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Dispersal

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 atintermediate and high altitudes (4500-13000 m). At high levels, atmospheric flow is laminar, but atlow levels, it is turbulent. Ash can be transported at 100-200 km/hr in the upper atmosphere. Onceparticles 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 ofgravity is balanced by the decelerating effect of air resistance. Beyond that point, the velocityremains constant, and ash can remained suspended until the wind velocity drops below theparticle's settling velocity, Vt, defined by Stokes Law:

Vt = {(8Rρsg)/(3Cρv)}0.5

where R is the particle radius, ρs is the particle density, C is the drag coefficient of the particle, and

ρv is the density of the transmitting medium.

In general, greater amounts of tephra fall out of the ash cloud near the vent, so that airfalldeposits typically thin away from the vent. However, secondary thickness maxima may occurdownwind. Airfall deposits typically have a circular or regular to irregular, fan-shaped distributionwith 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 thedispersal 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 istermed mantle bedding, as they typically "mantle" or drape over the underlying topography exceptwhere it is rugged. Bedding planes are distinct where deposition is on weathered or erosionalsurfaces, or different rock types. They may be gradational if deposition is slow by smallincrements so that bioturbation, wind reworking, and other soil-forming processes dominate. Thefabric in beds is commonly isotropic because elongate fragments are uncommon, with theexception of platy minerals and glass shards.

Airfall deposits are generally well-bedded and well-sorted, with bedding becoming morepronounced as sorting increases, and with size and sorting parameters varying geometrically withdistance from the source within single layers. Inman parameters (sφ) are commonly 1.0 to 2.0within both relatively coarse-grained as well as fine-grained tephra. Median particle diameters(Mdφ) are commonly -1.0 to -3.0 (2 to 8 mm) or smaller (phi values) close to the source, butfarther away may vary from 0.0 (1 mm) to 3.0 (1/8 mm) or more The sorting of airfall depositsdepends on:

(a) the distance from the vent;(b) variations in the strength and duration of eruptions;

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(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 reversegrading may occur in waterlain pumiceous deposits, and cross-bedding may result due to shifts inwind strength and direction. Differences in the proportions and densities of lithic, crystal and vitricconstituents produce both lateral and vertical variations in the size and nature of particles and natureof the deposits.

Vertical variations usually show increasingly basic compositions, often reflecting thetapping of compositionally zoned magma. Lateral variations reflect differences in the settlingvelocities of particles. Most ash deposits become more silica-rich with distance from the vent, asdifferent 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 thatloses its coherence as it approaches the ground surface. During phreatomagmatic eruptions, themagma is chilled and fractured on contact with ground or surface waters, resulting in violent steameruptions. In low-viscosity magmas droplet shape is, in part controlled by surface tension, byacceleration of the droplets after they leave the vent, and by air friction. The ash particles consist ofmostly 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 viscositymagmas, the morphology of ash particles is controlled primarily by vesicle density and shape, thevitric fraction generally consisting of very angular pumice fragments and thin vesicle walls brokenfrom pumice fragments during or after eruption. The morphology of lithic fragments is dependenton the texture and fracture pattern of the rock type broken up during the eruption. The morphologyof ash particles from phreatomagmatic eruptions is controlled by the thermal stresses within thechilled magma, which result in fragmentation of the glass to small blocky or pyramidal glassparticles. Vesicle density and shape play a minor role in determining the morphology ofphreatomagmatic ash particles.

P y r o c l a s t i c F l o w A n d S u r g e D e p o s i t sP y r o c l a s t i c F l o w A n d S u r g e D e p o s i t s

Pyroclastic flow and surge deposits represent the movement of large volumes of pyroclasticmaterial with the general behavior of lava flows. They act as heavy fluids controlled in theirmovement by gravity and the topography of the underlying land surface. The flows may be relatedto dome collapse, to explosive activity at the crater of composite volcanoes, or to discharge fromfissures. The deposits, which have variously been termed ashflow tuffs or ignimbrites, are diverseand reflect different types of eruptions and depositional regimes. From the perspective of flowmechanic, 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 andbetter bedded than pyroclastic flow deposits.

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Pyroclastic flows originate in different tectonic and volcanic settings and have vastly differentvolumes. Eruptions producing pyroclastic flow deposits on the order of 0.001 to 1.0 km3 arefrom 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 km3 originate from larger stratovolcanoes:

(a) 1883 Krakatoa, Java(b) Mount Mazama

Volumes of 100-1000 km3 are associated with the formation of large caldera which develop as aconsequence of eruption of such large volumes and not necessarily at the site of a pre-existingvolcano:

(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 km3).

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 theflows originate:

(a) Small-volume flows produced by dome collapse or explosions associatedwith dome formation commonly contain abundant poorly vesiculatedproducts of the domes, although some dominantly pumice flows also occur.

(b) Intermediate- to large-volume flows are usually composed entirely of highlyvesiculated materials derived from the rapid vesiculation of magma.

Relationship to Topography

Pyroclastic flows may completely drain from upper slopes and only be preserved in thelower parts of valleys, thereby becoming initially thicker away from the source. Pyroclastic flowsmay be confined to valleys. On the upper slopes of volcanoes, pyroclastic flows drain down valleycenters, leaving levees or high-water marks and larger rock fragments on both sides of a valley, oralong the outer edge of a sinuous channel because of the momentum of flow. Beyond mountainslopes, pyroclastic flows spread out in fan-like lobes. Widespread sheet-like ignimbrite layersassociated with large calderas and other volcano-tectonic depressions commonly have sufficientvolume and thickness to smooth out underlying topography. They:

(a) thicken and thin according to topographic irregularities underneath;

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(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 ofa valley, and their deposits may mantle topography similar to fallout tephra, but unlike fallouttephra, 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 flowunit is a depositional unit that represents a single pyroclastic flow deposited in one lobe. Thethickness of individual flow units can vary from a few centimeters to tens of meters, and the lobesmay follow one another within minutes or hours. The boundaries between flow units are markedby:

(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 singlecooling unit. A simple cooling unit forms when a single flow or successive flows cool as a unitwith no sharp changes in the temperature gradient. A compound cooling unit forms when there isan interruption in temperature that disturbs the continuous cooling unit zonation of successive hotflows.

Cooling from emplacement- to ambient temperature may take tens of years, depending onthe thickness of the deposit and the emplacement temperature. Thus, many ash flow deposits maremapped as cooling units, even though they composed of several flow units. A cooling unit ismarked by a more or less symmetric patterns of zones of rock differing in degree of welding andthus density resulting from different cooling regimes. In young deposits:

(a) The top and bottom parts of cooling units are commonly composed of friableunwelded pyroclastic material. The basal layer is unwelded because it coolsquickly against the cold rock basement, and the top because of relatively rapidheat 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 completecrystallization (primary devitrification) of hot and compacting glassypyroclasts occurs in the interior of thicker cooling units. Such zones gradeinto poorly welded zones lithified by crystallization of high temperaturevapor-phase crystals, typically silica polymorphs (tridymite, cristobalite) andalkali-feldspar.

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(c) The most obvious breaks between cooling units is the presence of anerosional surface, although breaks may also be indicated by reversals inporosity 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 magmaand the origin of the flows. Ash-flow tuff, by definition, is composed of more than 50 percent ofcomponents in the ash size range (<2 mm). These particles form a matrix into which varyingamounts of pumice lapilli or lithic lapilli are present.

(a) Glass Shards The most common ash-sized clasts are glass shards, usuallyaccompanied by smaller amounts of pumice particles. The ash and lapilli-sized pumice fragments are characterized by either spheroidal or longsubparallel tubular vesicles a few millimeters to micrometers in diameter.

(b) Crystals These are the next most common ash-size component. Crystals inignimbrites, which range from 0 to 50 percent abundances, are commonlybroken. Phenocrysts within accompanying comagmatic pumice lapilli orblocks, however, are largely non-fragmented, which indicates that breakageoccurs during eruption and transportation. Breakage may continue evenduring compaction. Crystals are generally more abundant in the matrix than inpumice lapilli and bombs, suggesting preferential concentration in the matrixrelative 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. Thereare three major sources for lithic fragments:

(a) slowly cooled and crystallized magma rinds from chambermargins;

(b)rocks from conduit walls; and, (c) rock fragments picked up along the path of the pyroclastic flow.

Charac ter i s t i c s o f Ash-F low Depos i t sCharac ter i s t i c s o f Ash-F low Depos i t s

Most unwelded pyroclastic flow deposits are poorly sorted and massive, but may showsubtle grading, alignment bedding or imbrication of oriented particles. In contrast, most pyroclasticsurge 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,

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(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. Pumicefragments may be inversely size graded and lithic fragments normally graded within the samehorizon because of their widely different densities. Slight differences in size of fragments indifferent layers may also give an irregular and indistinct stratification to the deposit.

Flat fragments within pyroclastic flow deposits may be strongly oriented parallel todepositional surfaces , or may become imbricated near the basal parts of the deposits due toshearing 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 poresor cavities called lithophysae. Pyroclastic flow deposits are also commonly characterized bynumerous degassing pipes or fumaroles. These degassing pipes may be recognized a oxidizedzones or as fines-depleted pipes in lower-temperature deposits. Minerals precipitated by fumarolesmay built fumarolic mounds and ridges, and are numerous where crystallization of an ash sheet ismost intense; the vapor-phase deposits involve minerals such as tridymite, alkali-feldspar, hematiteand 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 berelatively large features that stand 0.5 to 15 m above the flow surface and are up to 60 m indiameter; the ridges may be as long as 600 m. These zones typically represent zones of intensedevitrification, the late stages of which may lead to development of spherulites.

Textural Relationships

Most pyroclastic flow deposits are poorly sorted, having sorting values (sφ) that are greaterthan 2.0, and tend to decrease, as do median diameter values (Mdφ) with length of transport. Sizedistribution curves of different types of pyroclastic flow deposits tend to follow a Gaussiandistribution, which indicates that sorting takes place during eruption and transport. Poor sorting ischaracteristic throughout the length of a single pyroclastic flow sheet, but sorting varies verticallyat any single locality and tends to improve slightly with distance. Maximum size of both lithics andpumice decrease with distance in subaerial deposits.

In the textural analysis of pyroclastic flow deposits, it is important to know the relativeproportions of pumice, lithics, and crystals because their size distributions, sorting and otherparameters of these three subpopulations may differ for reasons other than sorting in the eruptioncolumn and during flow. Glass particles may diminish is size toward the margins and distal endsof the flow, due to breakage during transport. Lithic fragments could be derived from magmaticstoping, by fragmentation of the walls of the magma chamber and vent, or fragmentation of a plugdome 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 magmaand of breakage during explosive eruptions. Pumice has low mechanical strength and therefore

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may be reduced in size during eruption and flow, causing a preponderance of pumice dust in thefine-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 vitricparticles. Fragments with relatively low settling velocities are carried high into the atmosphere anddo not enter the flow, and others escape from the top of the flow while it is moving. Theseprocesses give rise to a distinctive type of fine-grained fallout tephra that has a crystal/vitric ratiosymmetrically lower than that from manually crushed pumice. Crystal enrichment may be greatestin the basal layers than in the middle of some ignimbrites, due to selective loss of glass shardsfrom 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-conservingmechanism. Mixing of hot pyroclastic material with cold air during flowage is minimal andrestricted to a thin surface of the flow. Hot pyroclastic flows may be nearly at magmatictemperatures during movement and shortly after emplacement. The emplacement temperature isdetermined 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 meterswithin 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°Cfarther away, with latter eruptions being hotter than earlier ones, and temperatures withinindividual flow deposits not decreasing substantially along their flow paths. The cooling rate isinitially 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 featuresof ignimbrites, e.g. the plastic deformation and welding together of glass shards. In compaction ofan ashflow deposit, two kinds of compaction are present:

(a) mechanical compaction: Mechanical compaction takes place as the result ofsimple loading without significant change in particle shape. Particles maintaintheir relative positions after coming to rest except that elongate particles tendto be rotated toward the horizontal. Pumice fragments generally maintain arandom orientation. Thus, the primary effect of mechanical compaction is toproduce reduced porosity:

(b) welding compaction: Welding compaction results from variable viscousdeformation of vitric fragments, from completely undeformed shard typical oflow-temperature fallout deposits to nearly homogeneous solid glass typical ofobsidian in which there are only ghost-like outlines of former shards in acontinuous glass matrix. The main control is the amount of time that

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temperatures remain above the threshold for welding (about 550°C). Thethreshold temperature depends on volatile content and chemical compositionof the glass. Complete welding may cause collapse of porous pumice to formdense, black, glassy discs, or fiamme, which are normally drawn out in aparallel structure or banding referred to as eutaxitic texture. The degree ofwelding 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 highertemperatures despite reduced thicknesses. Columnar jointing is common in moderately to highlywelded tuffs.

Structures Related to Temperature and Viscosity

Several structural features of welded to partly welded tuffs are related to the effects ofviscosity that develop during movement and deflation of a pyroclastic flow. Three main groups offeatures 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-depositionalcompaction of high-temperature pyroclastic flows causes flattening and welding of vitric shardsand pumice. Measurements of flattening ratio (F = length/width) of pumice fragments haverevealed a steady increase in the flattening from the top downward into the body of ash-flow sheetsdespite generally parallel changes in density.

Class i f i ca t ion and Nomenc la ture o f Pyroc la s t i c F lowsClas s i f i ca t ion and Nomenc la ture o f Pyroc la s t i c F lows

Classification and nomenclature of pyroclastic flows and their deposits have been thesubject of much confusion and debate. The wide range in physical properties of the eruptingmagmas, the different ways that flows originate, differences induced by transport processes, anddifferences in compaction and cooling structures and textures have given rise to a large array of

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names for the flows and their deposits. Traditionally, different kinds of pyroclastic flow have beennamed according to the volcano where first observed. Williams (1957), for example, arranged thedifferent types of flows in order of increasing gas content, increasing volume, and diminishingviscosity 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 ofthe flanks of domes and summit-spines or by breakup of the snouts andlevees 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 ofdomes. They are not buoyed up by exsolving gases, but rather slide. Thechaotic, unsorted, and unstratified deposits are generally less extensivethan those caused by explosions; they differ also in containing lesspumice, 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 andduring the rise of volcanic domes. Some Peléean flows are caused bylow-angle blasts. The ejecta vary from almost wholly lithic to almostwholly pumiceous. Largest, most gas-rich, and most destructiveavalanches usually occur during the initial phases of the growth ofdomes; these early flows are composed entirely of fresh effervescingmagma in the form of ash; blocky lithic debris is quite subordinate.Subsequent avalanches issue from the flanks of growing domes, and thedeposits being much coarser and include abundant lithic blocks. Depositsvary greatly in form and texture, are almost wholly confined totopographic depressions. They are heterogeneous, unstratified andunsorted. They are made up of angular blocks, torn from the solidcarapace of the parent dome, subangular, still effervescing bombs fromthe plastic interior, and sand- to dust-sized debris, some derived fromeffervescing magma and bursting bombs and some from larger fragmentsbroken in transit. Some blocks are of enormous size: up to 8 by 5 metersacross. Almost all fragments, particularly those in which the proportionof fresh magma was large at the time of eruption, are characterized by ahigh degree of porosity. Layers and lenses of sand- and dust-sized ashare commonly interbedded with the coarser debris. Peléean deposits arecharacteristically monolithologic and consist wholly of accessory andjuvenile debris of uniform composition. Accidental lithic fragments arerare. The deposits are not welded, though they may be indurated bycompaction.

2. Flows from summit-craters:

(a) St. Vincent type. Flows produced by backfall of ejecta from the marginsof 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 thenaccelerates as it interacts with the atmosphere; a lighter upper part rises

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because it has a higher temperature and hence a lower density than theatmosphere. When the fall-out of ejecta and heating of entrapped air failto reduce the effective density of the eruption-column below that of theatmosphere, the mixture of gas and solids falls back toward the vent andspreads outward as a glowing avalanche. Reduction of the gas-content ofan erupting magma may produce a change from pumice showers topumice 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 isexplainable by the mechanism of collapse of the mixture of hot gas andsolids; the higher particles reach in the eruption column, the greater thedistance of pyroclastic flow after collapse. The deposits are unstratified orpoorly stratified, poorly sorted, and unwelded, and are characterized bytheir fine texture and richness in crystals. Bombs and blocks make uponly 3 to 5 percent of the total volume; even lapilli constitute only a smallproportion. More than 90 percent of the deposits are of sand-size.Crystals make up 45 percent of most bombs, but they make up 73 percentof the ash. The concentration of crystals relative to vitric ash in depositsonly 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 upwellingfrom summit-vents and circumferential fractures on composite cones;eruptions commonly result in the formation of a caldera. Pumiceouspyroclastic flows erupted from summit-vents of large compositevolcanoes during late stages in their history. They are almost invariablypreceded by pumice falls, and usually involve magmas that are more fluidthan those of the Peléean, St. Vincent, or Asama types. Consequently,there is much more pumiceous material among the deposits andcorrespondingly less lithic debris. Pumice flows follow pumice falls asgas pressure diminishes. The pumice falls are well-sorted, well-stratified,and diminish in thickness away from the source. The deposits of thepumice flows are poorly sorted, and their stratification is irregular andindistinct. Except near the top, where coarse fragments are concentrated,they consist chiefly of sand-sized pumice. The compositions of thepumice-fall and pumice-flow deposits are essentially identical. Thesurfaces of the flows, particularly near their edges, are characterized bysubparallel ridges, up to a meter high, composed of large pumice lumpsthat 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 typeand those of St. Vincent and Krakatoan types. Initial discharge of gas-rich magma produces a pumice fall followed by two pyroclastic flows ofdiminishing gas-content. The pyroclastic flows, which display featurestransitional between those of Peléean and Krakatoan flows, are issuedfrom a summit crater as the magma foams over the crater rim and sweepsdownslope. The vesicularity of the ejecta decreases during the eruption.The earliest flows consist of material plastic enough to anneal; even at themargins of the flow, where the deposits are less than 30 cm thick, theyare compact and crudely jointed. No fragments, however, show

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flattening. Later pyroclastic flow involve more viscous magma, and manyof the ejected fragments are solid; these flows do not spread as widely asthe early flows, and tend to be confined to narrow channels. The depositsconsist of dense or slightly vesicular blocks and bombs and little juvenileash. Some blocks measured up to 30 m across None of the later flows arewelded or indurated by compaction.

3. Flows discharged from fissures:

(a) Valley of Ten Thousand Smokes type. Eruptions from one or more shortlinear fissures, the location and trend of which are unrelated to a cone orcrater. These flows consist predominantly of sand- and dust-sizepumiceous particles mingled with lapilli, and are almost completelyunstratified. Bombs and lithic blocks are rare. The deposits are almostcompletely unsorted. Lenses of cross-bedded, fluviatile pumice separatesome of the flows; these are laid down by floods when rivers burst dams.Most deposits are only weakly indurated, but some are slightly weldedand show columnar jointing. The flow are not characterized by distortionof glass shards or flattening of pumice lumps.

(b) Valles type. Eruptions of siliceous pumice from arcuate fissures formedby regional arching of the roofs of large bodies of rising magma; volumesof ejecta are usually so great that the roofs of the magma chamberscollapse along the arcuate fissures to produce calderas. These voluminouspyroclastic flows are discharged from arcuate fissures that develop whereeither there has been little or no volcanism for long periods or wherelong-continued eruptions have produced thick volcanic deposits. Alloriginate above large bodies of siliceous magma that dome their roofs.Once the fissures open, foaming magma escapes in huge volumes, firstproducing airfall deposits, and then ignimbrites. The volume of air-fallpumice is much less than that of the ignimbrites and the latter arecommonly welded. So much magma is expelled that the roof of themagma chamber usually collapses along the arcuate fissures.

Alternatively, Wright and others (1980) named the flows and their deposits according to fieldcriteria such as relative amounts of poorly vesiculated blocks, and ash (see Tables 12-1 to 12-6 inCas 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.

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V I I . L a h a r i c D e p o s i t sV I I . L a h a r i c D e p o s i t s

General FeaturesGeneral Features

Lahar is Indonesian for volcanic breccia transported by water, synonymous volcanic debrisflow, a mass of flowing volcanic debris intimately mixed with water. The term refers both to theflowing debris-water mixture, and the deposit thus formed. Many lahars are associated withstratovolcanoes of which they may comprise significant volumes of bulk. Most lahars arerelatively limited in extent, and occur in valleys or on alluvial aprons or lowland areas immediatelysurrounding volcanoes.

Many lahars are initiated directly by volcanic eruption, whereas others originate in wayssimilar to nonvolcanic debris flows. Once flow begins their fluid characteristics appear to besimilar or identical. The flows are non-Newtonian fluids that have a yield strength, behaving likeplastic material similar to wet concrete, have a high bulk density, and exhibit the property ofstrength which greatly influences the final textures and structures of the deposit. The Newtonianproperties of water (i.e. lacking in yield strength) begin to be modified by particle interferencewhen 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 withlaminar motion. As velocity decreases, the entire flow stops rather abruptly, after which waterseparates from the granular material by percolation or evaporation. On steep slopes, velocities maybe rapid enough to keep the entire mass in motion, but as slope decreases, the mass congealsunless it is thick enough to maintain a high shear stress at the base of the flow. As the gradientdecreases, velocities decrease, and the flow thins, shear stresses increase until the flow congeals toits 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 leavethin deposits on steep slopes and in the headwaters of valleys, but become thicker in valleybottoms and form fans that coalesce or else form broad individual lobes in lowland areas on verylow 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 aconstriction momentarily causes a large debris flow to pond up to several tens of meters above thevalley bottom and then drain away.

Lahars vary greatly in thickness. They tend to maintain a relatively constant averagethickness on relatively low slopes but locally vary depending on the configuration of underlyingtopography. Lahars come to rest with steep sloping lobate fronts. Most lahars are probably lessthan 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 anddepressions interpreted to be caused by differential compaction over an irregular underlyingsurface. The form, shape, and size of irregularities depends on the viscous properties of the flowsand the number and characteristics of multiple lobes.

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Basal Contact of Lahars

Although lahars may be very thick and carry large boulders, they commonly do not erodethe surfaces on which they flow except on very steep slopes. Lahars can pick up loose materialsfrom surfaces on steep slopes or where local turbulence develops within the flow owing to highlyirregular channels. Some Pleistocene lahars in the southern part of the Puget Sound lowland havetraveled 60 to 80 km from their source without picking up appreciable debris from the surface onwhich 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 oravalanching 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 daciticcomposition mixed with ash-sized minerals and lithic particles. Many lahar deposits containcharred wood, indicating that they were initiated as hot pyroclastic flows then cooled down duringtransport.

Grain-Size Distribution

Particles carried by lahars range from clay- to boulder-size, but the percentages of each sizefraction vary enormously from deposit to deposit and within a single deposit. In general, lahars arecoarser-grained and more poorly sorted than pyroclastic flow deposits. Grain-size parametersshow the obvious fact that lahars have a wide range in grain size and are coarse-grained and poorlysorted.

The presence of large boulders, commonly exceeding l m in diameter, is one of the mostcharacteristic features of lahars except perhaps, in their terminal zone. Large fragmentsprogressively decreased in number and size away from the source, although the finer constituent(matrix) may not show corresponding changes. Erratic fluctuations in median diameter areattributed to the longitudinal inhomogeneity of the flow caused by deposition from individualdebris 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 irregularbut slightly more concentrated arrangement of large fragments a short distance above the base; suchlayers are reversely graded. The large fragments in a lahar rarely rest directly upon the depositionalsurface. Some workers have suggested that large boulders re suspended by turbulence, it hasgenerally 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 thematrix. Differences in grading, whether it be absent, weakly or strongly developed, normal orreverse, appear to be related to the relative concentration of solids and fluids; the lower theconcentration of solids, the more likely normal grading develops because viscosity, density, andstrength of the fluid are less able to support large dense particles as velocity decreases. Whereconcentration values and viscosity, density, and strength are high, reverse grading develops,especially if the density of fragments is relatively low.

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Fabric

The fabric of lahars is commonly regarded as isotropic, but in some lahars disc-shapedpebbles and uncharred twigs and tree trunks concentrated low in the central parts of the deposit areoriented subparallel to base. The development or lack of clast fabric in the flows depends on themechanism of movement and deposition. Matrix strength in the flows may produce a rigid plugwhere shear stress is below the yield threshold, and this plug rides on a zone of laminar flowwithin which the shear stress is greater than the yield threshold. Flow stops when the plug expandsto the base of the flow at the expense of the zone of laminar flow

Orig in o f LaharsOr ig in o f Lahars

The mechanisms of lahar formation can be grouped into three major categories:

l. Direct and immediate result of eruptions: eruptions through lakes, snow orice; heavy rains falling during or immediately after an eruption; flowage ofpyroclastic 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 drainageof lakes or the avalanching of loose debris or altered rock.

3. Not related in any way to contemporaneous volcanic activity: mobilizationof loose tephra by heavy rain or meltwater; collapse of unstable slopes;bursting of dams due to overloading; sudden collapse of frozen groundduring the spring thaw.

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V I I I . S t r u c t u r e s B u i l t A r o u n d V o l c a n i c V e n t sV I I I . S t r u c t u r e s B u i l t A r o u n d V o l c a n i c V e n t s

The accumulation of erupted material around a vent forms a hill or mound. The form of thisstructure depends partly on the form of the vent and partly on the violence of eruptions. It alsodepends 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 builtby single eruptions may vary from a few meters to several hundred meters in height, and a fewmeters to more than a kilometer in width. Repeated eruptions build larger structures.

Cinder ConesCinder 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 conesresemble cinder cones in form, internal structure, and size range, but they differ in that they arecomposed of light-colored, more siliceous ejecta, such as rhyolite or dacite. Cinder cones typicallyrepresent rapid growth. For example, during its first 24 hours, Paricutin (Mexico) rose between 30and 50 m; when it was only a week old, it was 140 m high. After a year, it was 325 m high andhad a volume of approximately 0.2 km3. Over 90 percent of the fragmental ejecta were dischargedduring the first year.

External Form

Most cones are essentially symmetrical with slopes of 25° to 40°, and saucer-, bowl-, orfunnel-shaped summit-craters. They may overlap or coalesce or even grade into ramparts along acommon fissure. The forms of craters are a function of the nature of explosive eruptions and themechanical properties of ejecta. In general, a shallow explosion focus provide a broad, shallowcrater. The initial form is normally funnel-shaped, but slumping of walls, subsidence of materialinto 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 maximumangle of repose of the ejecta. If most eruptions are weak, the slopes of the cones tend to beconcave; 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 ofpyroclastic 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.

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Breached 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 arerafted away.

Internal Structure

Ejecta that make up Cinder cones are well stratified, with layers of different coarsenessranging from bombs to ash and consisting mostly of lapilli. Frequent and repeated size changesreflect variations in the strength of the explosions. The ejecta display little gravity sorting; only atgreater distances from the vents does tephra show graded bedding. Lava may be interbedded withthe scoria, but it is more typical for lava to issue from a vent (bocca) at or close to the base. Ifpresent, dikes are irregular in trend and width and may intersect in a complex manner. Most ejectaare coarsest and thickest close to the vent. Ejecta from violent eruptions diminish outward inthickness and size less rapidly than do the ejecta laid down by weak eruptions.

Maar VolcanoesMaar Volcanoes

Maar volcanoes are low volcanic cones with bowl-shaped craters that are wide relative torim height. They were originally recognized as small subcircular crater lakes, the term beingderived from the Latin "mare" for sea. The various kinds of maar volcanoes are:

1. Maar (sensu stricto): Volcanic crater cut into country rock below generalground level and possessing a low rim composed of coarse- to fine-grainedtephra. They range from about 100 to 3000 m wide, about 10 to more than 50m deep and have a rim height of from a few meters to nearly l00 m abovegeneral ground level.

2. Tuff ring: Large volcanic crater at or above general ground level surround by arim 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, andare 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 craterstands high and another side low.

Most maars result from hydroclastic eruptions; wide craters develop from shallowexplosions, subsidence or a combination of both. In groups of nearly synchronous eruptivecenters, those erupting on high ground form spatter or cinder cone whereas associated eruptioncenters in valleys, depressions, or alluvial gravels in coastal regions form maars, tuff rings or tuffcones. 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 glassresulted from thermal shock and (steam) explosions, and that the vapor and steam phase in theeruption column was partly or largely vapor from external water.

Tuff cones and tuff rings are distinct landforms that result from slightly different types ofhydroclastic activity an represent a "continuum" of landforms from cinder cones to pillow lavarelated to environments of eruption and mechanical energy of eruptions. Tuff rings evolve througha stage of explosion breccia emplacement to a stage dominated by base surges which deposit thinlybedded layers. Tuff cones may be built when continuing activity evolves into a third stage,

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characterized by rock emplaced by poorly inflated base surges and ballistic fallout. The differencesare related to water: melt ratios; fragmentation of melt attains maximum explosive energy when thewater:melt ratio is about 0.5 for basaltic compositions, whereas initial ("vent-coring") eruptionswith 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 astuff rings. Higher ratios produce "wetter" and denser eruption columns giving rise to poorlyexpanded surges, hence dominantly massive beds and tuff cones. The rates of magma and waterinflux controls the process, and such "cycles" may be interrupted, reversed or alternate. Mostcommonly, tuff cones may have craters filled or partly filled with lava, and agglutinated spatter andcinders. In some volcanic fields, scoria cones contain deposits of phreatomagmatic origincommonly developed during their initial eruptive stages.

Lit tora l ConesLi t tora l Cones

Littoral cones are mounds of hyaloclastic debris constructed by hydroclastic explosions atthe point where lava enters the sea, and represent craters that lack feeding vents connected tosubsurface magma supplies ("rootless") and form where lava or pyroclastic flows move over smallponds of water, swamps, springs or streams. The cones commonly occur as crescent-shapedridges, breached by the source lava or more rarely as complete cones with craters occurring abovelava tubes. Explosion centers are near or at the shore line, therefore about half of the radiallyexploded material falls into the sea, leaving a half-cone on land. A typical littoral cone ischaracterized 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 bedsranging 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 iscommonly no more than 5% of the total ash content, and is composed of sideromelane,microcrystalline basalt and broken phenocrysts. Some layers contain accretionary lapilli andbedding sags, suggestive of abundant water vapor in the explosion clouds.

Shie ld Vo lcanoesSh ie ld Vo lcanoes

Shields are mainly confined to volcanoes produced by rapid accumulation of fluid basalticlavas. Three principal types of shields can be distinguished: the Icelandic, Hawaiian, andGalapagos types.

Icelandic Shields

These simple and most symmetrical shield volcanoes are formed entirely or almost entirelyby effusive eruptions from central summit-vents. Icelandic shields range in height between 50 and1,000 m, averaging 350 m. The angles of slope are unusually small, varying from 1° to 5°, butexceptionally may be as steep as 10°. The summit-craters or sinks of Icelandic shield volcanoes areapproximately circular, and most measure less than l km across. They have raised rims built ofspatter from lava fountains and by repeated overflow from lava lakes. Most summit-craters areessentially cylindrical pits with flat floors that may contain smaller collapse pits. A few collapse

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pits may also be found on the flanks of the shields, but they are distributed at random. There arefew radial fissures and lines of parasitic cones on the flanks, and concentric fissures around thesummit-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 smalllava-tubes.

Hawaiian Shields

These shield consists of three superposed units. The lowest and largest is made up ofpillow basalts and other products of submarine eruptions from deep-water vents. These areoverlain by hyaloclastites formed by eruptions in shallow water and by subaerial lavas that enteredthe sea. The topmost unit consists of thin, subaerial flows. Profiles of Hawaiian shields duringmature stages of growth tend to be slightly convex in their upper parts, uniform in their middleparts, and slightly concave in their lower parts. The upper slopes are only about 45 to 65 m perkm, whereas the middle slopes are three to four times as steep (ca. 10° ), and around the base theydiminish to 2° or 3°. During the mature, shield-building stages of growth, basalts are discharged atintervals of a few years from vents near the summit and along radial rift zones. Virtually all of thelavas are of the pahoehoe type near their source, but downslope many change to aa. As the shieldsgrow, 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 laststages of development.

Following the mature stage of growth, the volcanoes pass into a declining stage when theshields are buried beneath steeper-sided structures built of thicker, and more varied types of lavaand pyroclastic debris, and by clusters of parasitic scoria cones and trachyte domes. During theirinitial stages of growth, Hawaiian shields may resemble Icelandic shields in that they are built byrepeated outflows from central summit-vents. Subsequently, the summit activity is accompanied byan increasing number of eruptions from rift zones on the flanks. Mature Hawaiian shields aredistinguished by prominent rift zones that converge at the summit-calderas. Main rift zones resultfrom 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 thelowest lava-vents, and the maximum outflow is not always from the lowest vents, but from thewidest parts of the fissures. Locally, short en echelon eruptive fissures progress up- slope ratherthan downslope. Eruptive fissures may cross the floors of summit-calderas and collapse pits. Riftzones, 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 orsteeply inclined dikes.

Galapagos Shields

Shield profiles in the mature stage of growth resemble those of giant tortoises or overturnedsoup bowls with deeply indented tops. Visible parts measure between 15 and 30 km across at sealevel, and they rise to heights of approximately 1,100 to 1,700 m. The slopes of some are less than20°, but the middle slopes reach angles of 15° to 35°, flattening rapidly near the base and just asrapidly to wide benches surrounding the summit-calderas. Some fissures are gaping cracks thatshow 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 theflanks of the shields, where they are accompanied by chains of scoria-cones, generally breached ontheir 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

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only partly by eruptions from radial fissures. Pyroclastic ejecta are present only during thedeclining stages of growth. Nordlie has presented a model for the origin of the Galapagos shields.

Compos i te ConesCompos i te Cones

Cones built partly of lava flows and partly of fragmental ejecta have long been referred toas 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 placemore and more from radial fissures far down the flanks, while explosive eruptions may continuefrom the summit. Radial fissures result from the tumescence of cones and from the increasinghydrostatic 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 thetypical concave profiles. Coarse ejecta blown from summit-vents accumulate close to their sourceand their deposits generally have high angles of repose, whereas fine ejecta, which accumulatechiefly 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 theupper slopes and depositing it around the base. In general, the older a cone and the longer it hasbeen extinct, the more pronounced is the concavity of its slopes. The shapes of large compositecones 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 formsintermediate between those of domes and shield volcanoes, whereas others, composed almostwholly of pyroclastic ejecta, tend to have symmetrical forms with uniform slopes. Compositevolcanoes have a conical form only if built mainly or exclusively by eruptions from a central, moreor less cylindrical conduit. Those built by eruptions from elongate fissures or subparallel fissureshave forms that resemble those of overturned canoes. Elongate forms may also result fromprogressive migration of the main conduit along a fissure system.

The basal parts of volcanoes that rise from a shallow sea floor differ from those builtwholly on land in that they contain much more fragmental material, both volcaniclastic debris andsediments, and are intruded by many flat-lying sills.

Internal structure

Dissection reveals essentially conformable, outward dipping layers of lava and pyroclasticdebris. In multiple-vent cones, there may be many angular unconformities, owing to overlap oflavas and pyroclastic debris from different vents. Dikes, sills, and central plugs are commonlyexposed by deep erosion. The dikes may be irregular in trend and thickness, but most areapproximately radial. Swarms of subparallel dikes are exceptional; so are ring dikes and cone-sheets. Sills are commonly mistaken for lava flows.

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Fillings of the central conduits, usually referred to as plugs, necks, or pipes, vary greatly insize and shape, as well as in composition, structure, and texture. Few plugs exceed a few hundredmeters in diameter, though some measure more than a kilometer across. They may pinch andswell, but most merge downward into stocks. Crudely cylindrical shapes are typical, but manyplugs 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 orcoalescing structures built by discharge of magmas of different composition and by eruptions ofdifferent kinds. Such volcanoes are compound and grouped genetically into normal, recurrent, andinverse types:

1. Normal: Cones in which successively younger, generally smaller forms arebuilt by eruption of increasingly differentiated magmas. Many large “normal”composite cones are composed principally of basalt and andesite while youngerproducts erupted on the flanks form domes of dacite and rhyolite or basalticcinder 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 ofwhich collapse to form calderas as a result of large explosive eruptions of moresiliceous magma. New composite cone then develop within the caldera andfollow a similar evolutionary sequence.

3. Inverse: Cones in which the normal differentiation series from less to moresiliceous magmas is reversed.

Parasitic Cones

The growth of parasitic cones on the flanks of large composite volcanoes is a sign of oldage. Not uncommonly, these cones develop at successively lower levels as the volcanoes approachextinct. Usually, they are made up of more basic and more siliceous differentiates. Parasites maybe concentrated along lines or belts that reflect structural trends in the subvolcanic basement, or in acrudely concentric arrangement. The concentric rings may reflect cone-sheets or ring dikes atdepth. A crudely radial arrangement of parasitic cones and domes is much more common. Thenumber of parasitic cones on most large composite cones is seldom more than ten or a dozen.

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I X . C r a t e r s , C a l d e r a s , a n d G r a b e n sI X . C r a t e r s , C a l d e r a s , a n d G r a b e n s

These volcanic features, which represent negative landforms or depressions, are of severalkinds and origins. Small to moderate sized depressions, when more or less circular in plan, arecalled craters. Larger quasi-circular depressions, which typically are many times greater than thatof any included vents, are referred to as calderas. The lower size limit of calderas is a 1.5-kilometerdiameter. 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 collapserelated types.

Explos ion CratersExplos ion Craters

These are normally formed at the summit of lava, spatter, cinder, and ash cones, and ofcomposite 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 oflava from the upper part of the conduit

Col lapse CratersCol lapse Craters

These are nearly circular craters that perforate the surface of a volcano without anysurrounding debris cone. They result from sinking in of part of the volcano, and range from a fewmeters to kilometers in diameter. They may be few meters to over 300 meters deep, but aregenerally roughly circular in plan. The wall of the crater are nearly vertical during early stages ofgrowth, but become more sloped with erosion.

CalderasCalderas

In form and origin, calderas resemble large pit craters, differing only in size. There floor isbroad and generally flat. Curved faults typically lie outside the main caldera and are more or lessparallel with the caldera walls. These walls are steep cliffs, formed by faulting, with banks of talusat the base formed by fragments falling from the cliffs. Outward-dipping lava flows in the calderawalls are typically cut off abruptly.

There is generally no question of a collapse origin because only small amounts ofpyroclastic material are present. The history of a caldera consists of a series of collapses andrefillings. The entire floor may rise as a unit, with the entry of new magma from below. Hawaiiancalderas, for example, appear to have grown by gradual coalescence of a number of pit craters thatformed, one after another, on the summit of the shields. Fault scarps are more or less tangential tothe 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 oflava flows and pyroclastic cones at the top of the shield. The crater filling lavas can be easilyrecognized from the main shield lavas by:

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(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 5-6miles wide and more than 1000 meters deep. Formation of these calderas usually takes place late involcanic activity, and often follows a long pause in activity during which the cone may be deeplyeroded. Magmatic material ejected during the caldera-forming eruption is generally rhyolitic ordacitic in composition. Commonly ash flows appear to have been erupted from the same series ofarcuate fissures on which the circular summit block subsides to form the caldera. Where formed ona composite volcano, the eruptive fissures are sometimes part way down the slope of the cone, butin other instances, the distribution of tephra indicates ejections came from restricted vents at thesummit. The volume of material thrown out by explosions equals only a few percent of the calderavolume. It is clear that the caldera has not resulted from explosive decapitation of the mountain, butfrom subsidence along ring fractures. A huge volume of juvenile magma erupted indicates collapseprobably occurs as partial drainage of an underlying magma chamber removes support frombeneath the summit of the volcano. Such an interpretation implies the presence of a fairly largemagma reservoir.

Class i f i ca t ion o f Ca lderasClass i f i ca t ion o f Ca lderas

Williams (1957) suggested that calderas fall into one of the following types: Krakatoan;Katmai; Valles; Hawaiian; Galapagos; Masaya; and Atitlán. Williams and McBirney furthersuggested that these types can be divided into groups distinct not only in the nature of the magmaswith which they are associated, but also in geophysical properties:

(a) Group I - moderate to strong negative gravity anomalies resulting from deepinfill of light fragmental material (Krakatoan, Katmai, and Valles types); and,

(b) Group II - moderate to strong positive gravity anomalies, resulting from thepresence 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 compositevolcanoes following explosive eruptions of siliceous pumice from one or more vents or, in someinstances, from arcuate fissures on the flanks. The volume of ejecta is usually much less than 100km3.

Prior to formation of the caldera, the site may have been occupied by a cluster ofcoalescing, composite cones. Eruption follows a long period of quiet, and may be immediatelypreceded by up to 16 years seismic activity. The climactic eruption probably lasts no more than 24hours, but in that brief time, 2 to 6 km3 of pumice are discharged. The eruptive column may risemore 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 beinfluenced by inflow of groundwater that causes phreatomagmatic outbursts to form thin layers ofvitric ash interbedded with accretionary lapilli, and simultaneous heavy rains convert some pumicedeposits into lahars. Finally, lava may be issued at the foot of volcano. As a consequence of theseeruptions, the top of the volcano collapses to produce a caldera. After collapse, intra-caldera domesof andesite and dacite rise at intervals from the floor

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The shape of the caldera may be been determined by three principal controls: 1) arcuatebays forming scalloped margins that may indicate the collapse of individual cones into cupolasabove a common reservoir; 2) radial grabens extending outward from the caldera that may reflectextensions 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 newvolcanoes 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 resultspartly from explosions, but mainly occurs when magma is drained from the central conduit toreplace and mix with magma in a neighboring fissure system. Both magmas are discharged fromfissures on the flanks.

Valles Type

Valles type caldera foundering takes place along arcuate fractures independent ofpreexisting volcanoes as a consequence of simultaneously with a discharge of colossal volumes ofsiliceous pumice, usually more than 100 km3.

Calderas of this type, which include the world's largest, are produced by collapseaccompanying and following discharge of great volumes of tephra from arcuate fissures formed bythe rise of large subjacent bodies of siliceous magma. The caldera typically measures at least 20 by25 km across, and its walls enclose a circular moat partly filled by a ring of rhyolite domes thatmay surround a resurgent dome (up to 650 m high and 10 by 13 km wide) produced by uplift ofthe 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 rangingin composition from rhyolites to basalts. The total volume of ejecta is close to 200 km3.

Foundering of caldera may temporarily close the arcuate eruptive vents. Pumice flowssweep 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 fissuresnear its center and a lake may form on the floor. A group of late rhyolite domes are typically builtover ring-fractures within the surrounding moat.

Hawaiian Type

Hawaiian type calderas are formed by collapse of the tops of shield volcanoes during latestages of growth. Prior tumescence is followed by subterranean drainage of basic magma frombeneath 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 theshields has grown almost to its full height and while eruptions are still vigorous and frequent. Asthe calderas increase in width and depth, lavas pour from fissures cutting the floors and walls, andfrom rift zones on the flanks. Ultimately subsidence comes to an end, and the caldera-filling stagebegins. Eruptions are spaced at much longer intervals and are more explosive. Flows and tephraaccumulate 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 widetraversed by inward-facing fault scarps. Both are also closely associated with pit craters; the

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calderas grow in part by coalescence of adjoining pits. It has been suggested that the calderas resultfrom drainage of the central conduits into rift zones and by eruptions of lava far down the flanks ofthe shields. However, the volume of lava discharged during any eruption is generally much lessthan the volume of the summit-collapse and concurrent subsidence. The calderas typically collapsefollowing broad tumescence of the shield when the rift zones are distended by magma, whether ornot 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 thefloor sinks as a wedge-shaped block, indicating that subsidence of the summit takes place afterfractures 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 basalticshields but engulfment results from injection of magma and eruptions of lava from circumferentialfissures 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 radialfissures on the flanks. The calderas are up to 7 by 10 km wide and up to 845 m deep. The numberand 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 owetheir origin to periodic distentions of the shields by rising magma. Radial fissures on the flanks ofthe 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 shallowdepression occupying most of the central portion of a low inconspicuous shield; eruptions fromarcuate and radial fissures outside the caldera play no part, and nearly all the lavas are containedwithin the boundary scarps.

The cauldron occupies most of the central part of a low basaltic shield, and measures up to6 by 11 km across. Its outer slopes are gentle. No arcuate vents border the rim and no rift zonescut the flanks of the volcano. The walls, which range up to 150 m in height, consist of basalticlavas 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, theirdisappearance is followed by localized collapse. The crater floor may subside as much as 200 mduring a single eruptive cycle. The scalloped margins of the depression suggest a succession ofroughly cylindrical collapses resulting from periodic migration of magma underground. Explosiveeruption 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 butassociated with eruptions from volcanoes near the rim or from nearby fissures. The cauldron iscaused by intermittent collapse, possibly resulting from underground migration of magma to feedadjacent volcanoes. Fault scarps that enclose one side of the basin may be deeply eroded and from300 to 600 m high, whereas on the other side, may become fresher, steeper, smaller, and passunder the summit of a caldera-rim stratovolcano.

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CauldronsCauldrons

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 exceedsthat 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 founderinginto a shallow magma chamber contribute to the development of both. The principal difference isthat calderas are associated with a withdrawal of magma, while cauldrons are thought to resultfrom a passive foundering of the roof of a static or rising body of magma.

Volcano-Tec ton ic Depres s ionsVo lcano-Tec ton ic Depres s ions

Volcano-tectonic depressions are bounded by faults of tectonic origin and may reachdimensions of tens or even hundreds of kilometers in length. These great down-faulted troughs areoften, though not always, formed on the crests of broad arches. For example, the Taupo Basin onthe North Island of New Zealand is more than 60 km long and 30 km wide and contains severallarge composite volcanoes. The basin formed by collapse of the crest of a broad arch, forming aseries of grabens, At about the same time, great masses of rhyolitic ash flows erupted fromfissures in the crest of the arch or from faults and fractures bounding the grabens. The compositevolcanoes only formed within the graben after collapse. Thus, these depressions are geneticallysimilar to Krakatoan and Valles-type calderas.

Resurgent CalderasResurgent Calderas

As already mentioned, volcanic activity commonly continues after caldera collapse. Thecaldera may eventually be completely filled and obliterated by later volcanic rocks. Renewal ofactivity in Valles-type calderas appears to be commonly associated by the up-bowing of the calderafloor, 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.

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X . C L A S S I F I C A T I O N O F V O L C A N I CX . C L A S S I F I C A T I O N O F V O L C A N I CE R U P T I O N SE R U P T I O N S

Volcanic eruptions have been classified on several different bases. The usual ones arebased on the form of the eruptive vent, the location of the eruptive vent, and the character of theeruption.

N a t u r e o f V e n tN a t u r e o f V e n t

The most generic classification is that based on the form and location of the vent, andinvolves:

(1) fissure eruptions - volcanic material issues from a crack or fissure. These eruptions tendto be typical of most oceanic volcanoes.

(2) central vent or pipe eruptions - volcanic material issues from a vent at the apex or centerof 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, whereasflank eruptions near the summit are referred to as subterminal eruptions. Terminal and subterminaleruptions produce no marked lowering of the top of the magma conduit in the conduit, but lateraland excentric eruptions generally result in a marked lowering of the magma column and an increasein gas activity at the summit. The later eruptions, which build parasitic or adventive cones, tend tobe wholly effusive in nature while explosive gas release occurs at the terminal vent.

S t y l e s o f E r u p t i v e A c t i v i t yS t y l e s o f E r u p t i v e A c t i v i t y

Volcanic eruptions have been classified on several different bases. The usual ones arebased on the form of the eruptive vent, the location of the eruptive vent, and the character of theeruption. In the latter case, the eruptive style is based on the physical nature of the magma, thecharacter of explosive activity, the nature of effusive activity, the nature of dominant ejecta, and thestructures built around the vent. Based on these criteria, six principal eruptive styles, most namedafter 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 lavaflows and less voluminous pyroclastic debris. Most eruptions start from fissures, commonly as aline of lava fountains, that ultimately coalesce to one or more central vents. Fragmental ejectanormally precede discharge of lava flows. Thin, fluid lava flows can gradually build up large broadshield 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 characteristicfeatures of these eruptions are:

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Physical 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 asecond to several hours in magma columns near the surface. Lava fountains are small andexceptional. Explosion clouds seldom rise more than 500 m, and are usually grey, with little orno lightning. Lava flows may follow explosive activity, and may continue uninterruptedly formonths or years. Volcanoes with this activity include Stromboli and Mount Etna, Sicily, Italy; andParicutin, 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 offresh, effervescing magma. Separation of a gas cloud from the avalanche produces a nuée ardentthat may move independently of the associated ash flow. Airfall ejecta are not widespread. Viscousmagma follows to form steep-sided domes and spines or short, thick flows, the flanks of whichmay collapse by gravity or internal explosions to produce hot block-and-ash flows. Volcanoes withthis 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 eruptionrate, voluminous, gas-rich eruptions that produce widely dispersed sheets of pyroclastic materialderived 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:

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Physical nature of magma: viscous; felsic (rhyolitic, trachytic, phonolitic, dacitic) becomingmore mafic during course of eruption.

Character of explosive activity: paroxysmal ejection of large volumes of ash, withaccompanying 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 thatproduces 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 widespreaddistribution. 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 explosionscaused 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 fragmentationoccurs chiefly due to thermal shock when the magma is quenched; particles are bound by fracturesand 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 ofStrombolian- or Hawaiian-type eruptions (into which they commonly grade as water is deniedaccess to the magma during the course of the eruption). The criteria are as follows:

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Strombolian/Hawaiian Surtseyan

Median diameter on or near cone 1/2 to > 16 mm usually 2 to < 1/8 mmSpatter common absentParticle shape achneliths common achneliths absentAccretionary lapilli absent commonImpact structures (bomb sags) rare or absent commonThickness of individual beds > 1 cm, most > 5 cm commonly < 1 cm to 1 mmAlteration always oxidized; never oxidized;

palagonite rare palagonite typical

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A p p e n d i c e sA p p e n d i c e s

A . P y r o c l a s t i c F a l l D e p o s i t sA . P y r o c l a s t i c F a l l D e p o s i t s

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 independentevidence for water flushing (e.g. accretionary lapilli, vesicles,water-splash microbedding). Fall deposits can be sufficiently hotto show primary welding when they accumulate near vent.

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B . P y r o c l a s t i c F l o w D e p o s i t sB . P y r o c l a s t i c F l o w D e p o s i t s

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 ofthermoremanent magnetization of contained clasts)

Exception - low-aspect ratio ignimbrites include a mantling layer whichpasses laterally into the valley-pond ignimbrite.

Note 1: ignimbrite can be defined as a pyroclastic flow deposit made mostlyof pumiceous material (pumice, shards)

Note 2: primary mudflows (lahars) resemble pyroclastic flow deposits butlack (d)

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C . P y r o c l a s t i c F l o w D e p o s i t C h a r a c t e r i s t i c sC . P y r o c l a s t i c F l o w D e p o s i t C h a r a c t e r i s t i c s

Deposit Descriptions

Ignimbrite Pumiceand Ash

Unsorted ash deposits containing variable amounts of pumice roundedsalic lapilli and blocks up to 1 m in diameter. In flow units pumicefragments can be reversely graded while the lithic clasts can shownormal grading; upgraded flows are as common. A fine-grained basallayer is found at the bottom of flow units. They locally contain fossilfumarole pipes and carbonized wood. The coarser smaller volumedeposits usually form valley infills while the larger volume depositsmay form large ignimbrite sheets. Locally, they may show one ormore zones of welding.

Scoria and Ash

Topographically controlled, unsorted ash deposits containing basalticto andesitic vesicular lapilli and scoriaceous ropy-surfaced clasts up to1 m in diameter. They may in some circumstances contain large non-vesicular cognate lithic clasts. Fine-grained basal layers are found atthe bottom of flow units. Fossil fumarole pipes and carbonized woodmay also be present. The presence of levees, channels and steep flowfronts indicate a high yield strength during transport of the movingpyroclastic flow.

Vesicular Andesiteand Ash

Topographically controlled, unsorted ash deposits containingintermediate vesicular (between pumice and non-vesicular juvenileclasts) 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 canexceed 5 m in diameter. The deposits are generally reversed graded.Fine-grained basal layers, fossil fumarole pipes and carbonized woodmay be present. Surface manifestations include the presence of levees,steep flow fronts and the presence of large surface blocks, all of whichindicate a high yield strength during transport of the movingpyroclastic flow.

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D . P y r o c l a s t i c S u r g e D e p o s i t s D . P y r o c l a s t i c S u r g e D e p o s i t s

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; depositsshow:

(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 bysedimentation from the pyroclastic flow.

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E . P y r o c l a s t i c S u r g e D e p o s i t C h a r a c t e r i s t i c sE . P y r o c l a s t i c S u r g e D e p o s i t C h a r a c t e r i s t i c s

Deposit Description

Base Surge

Stratified and laminated deposits containing juvenile vesiculatedfragments ranging from pumice to non-vesiculated cognate lithicclasts, ash and crystals with occasional accessory lithics (largerballistic ones may show bomb sags near vent) and depositsproduced in some phreatic eruptions which are composedtotally of accessory lithics. Juvenile fragments are usually lessthan 10 cm in diameter due to the high fragmentation caused bywater/magma interaction. Deposits show unidirectionalbedforms. Generally, they are associated with maar volcanoesand tuff rings. When basaltic in composition they are usuallyaltered to palagonite.

Ground Surge

Deposits are less than 1 m thick, composed of ash, juvenilevesiculated fragments, crystals and lithics in varying proportionsdepending on the constituents in the eruption column. Typicallyenriched in denser components (less well vesiculated juvenilefragments, crystals and lithics) compared to accompanyingpyroclastic flow. They show unidirectional bedforms;carbonized Generally wood and small fumarole pipes may bepresent.

Ash Cloud Surge

Stratified deposits found at the top of and as lateral equivalentsof flow units of pyroclastic flows. They show unidirectionalbedforms, pinch and swell structures and may occur as discreteseparated lenses. Grain size and proportions of all componentsdepend on the parent pyroclastic flow. Can contain smallfumarole pipes.

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