Geo.416 Volcanology

7/29/2019 Geo.416 Volcanology

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G E O . 4 1 6 VO L C A NO L O G YE O . 41 6 V O L C A N O L O G YI. Physical Natu r e Of Magma s . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4

Structural State of Silicate Melts . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 4Viscosity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 5

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

Yield Strength . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 8

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

II. Gener ation, Rise And Stor age 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 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 13

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

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

III. Er uptive 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 F lows . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23Volume . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23Length and Thickness . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 23


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

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

Morphology Of Lava Flows . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26

Pahoehoe Lavas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 26External 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. Pr oducts Of Volcan ic Explosions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33Terminology and Classification . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 33Origin . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 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. Lah ar ic Deposits . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47General Features . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47

Surface of Lahars . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 47

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

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

VIII. Stru ctur es Built Around Volcanic Vents . . . . . . . . . . . . . . . . . . . . . . . . . 50Cinder Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 50

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


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

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

Galapagos Shields . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 53Composite Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54External Form . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54Internal structure . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 54Growth Sequences . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55Parasitic (adventive) Cones . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 55

IX. Cr aters, Calder as, and Gr abens . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56Explosion Craters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56Collapse Craters . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 56Classification of Calderas . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57

Krakatoan Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 57

Katmai Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58Valles Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58Hawaiian Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 58Galapagos Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59Masaya Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59Atitln Type . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 59

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

X. Classification O f Volcanic Er upt ions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61Nature of Vent . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61Styles of Eruptive Activity . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61

Hawaiian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 61Strombolian Eruptions . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 62Pelean 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 . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 68

E. Pyroclastic Surge Deposit Characteristics . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . . 69


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G E O . 4 1 6 VO L C A NO L O G YE O . 4 1 6 V O L C A N O L O G YI . P h y s i ca l N a t u r e O f M a g m a s. P h y s i ca 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 ), 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.

S t r u c t u r a l S t a t e o f Si li ca t e M e lt st r u c t u r a l S t a t e o f S il ic a t e M e l t sModern concepts of silicate liquid structure are based on the Zachariasen Model. The atoms

are bonded by forces similar to those between atoms of crystals, but lack long range 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.

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V i s c o s i t yi s c o s i t yViscosity is a fluid's internal resistance to flow. It represents the ratio of shear

stress to rate of shear strain applied to a layer of thickness Z and permanentlydeformed in a direction x parallel to the stress. Mathematically, viscosity is expressedby:

s = so+

dm

dt

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; is the viscosity, expressed in unitscalled poises (dyne sec/cm2); dm/dtis the gradient of velocitydx/dt or strain rate over a distanceZnormal 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 < s o) 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 exhibitNewtonian behaviorwhen n equals 1.0 and soequals 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 they

contain 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.

Cont r ols 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 (poises)0:1:2 10101:1:2.5 282:1:3 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 1200C; 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 steadily

increases. For example:a) measured viscosity of a Mauna Loa flow increased 2-fold over a 12-mile-

distance from vent;b) measured viscosity of a small flow from Mt. Etna increased 375-fold in a

distance of about 1500 feet.

The decrease in viscosity can be attributed to an increase in distance between cations and anions,and therefore, a decrease in Si-O bond strength.

Time

At temperatures below the beginning of crystallization, viscosity also increases with time.

If magma is undisturbed at a constant temperature, its viscosity may continue to increase for 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.

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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. At

constant 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) dry (poises) wet (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, in

melts 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:

= o(1 - RC)-2.5

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where is the effective viscosity of a magmatic liquid, C is the volume fraction of suspendedsolids; o 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.

Y ie ld S t r e n g t hi el d S t r e n g t hMost magmas have an appreciable yield strength, which shows a marked increase below

their liquidus temperature. As yield strength increases, the stress required to initiate and sustainflow becomes greater, and the magma's apparent or effective viscosity is also increased.

S pec i f i c Hea tpec i f i c Hea tThe specific heat (Cp) of magma, which is the heat required to change the temperature of

the liquid 1 degree Celsius, is typically about 0.3 cal. gm-1. The specific heat contrasts greatly 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.

T h e r m a l C o n d u c t i v i t yh e r m a l C o n d u c t i v i t yIgneous rocks and liquids are poor conductors of heat. Thermal conductivity depends on

two heat transfer mechanisms:

(1) ordinary lattice or phonon conduction; and,(2) radiative or photon conduction.

The former declines and the latter increases as temperature increases and the melt 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 to1200C, 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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D e n s i t ye n s i t yMagma 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 cm2dyne-1.

E l e ct r i c a l C o n d u c t i v it yl ec t r i c a l C o n d u c t i v i t yElectrical conductivity, which is low in pure silica melts, increases with increasing

abundance 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 oc i t i e se i s m i c W a v e V e lo c i t i e sCompressional or P-wave velocities are about 6 km sec-1 up to the melting range, then

decrease abruptly to 2.5 km sec-1 at higher temperatures. Shear or S-wave velocities are about 2-3km 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 . G e n e r a t i on , R i s e An d S t o r a g e O f M a g m aThe subsurface processes by which magmas are generated and rise toward the surface are

extremely complex. Before examining these processes, it is worthwhile to review what is knownconcerning the Earth's interior.

N a t u r e of C r u s t a n d U p p e r M a n t l ea t u r e o f C r u s t a n d Up p e r M a n t l eMost of what is known concerning the Earth's interior comes from geophysical

measurements, and concerns:

(a) seismic wave velocities;(b) temperature;(c) density distributions;(d) heat flow; and,(e) mechanical properties.

Seismic velocities increase with depth within the Earth, but show abrupt changes at 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 and the elastic properties (bulk modulusK and rigidity or shear modulus ) by the following expressions:

Vp = { [K + (4/3)]}Vs = (/)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.

H e a t S o u r c ese a t S o u r c esExistence of magma indicates that at some depth beneath the Earth's surface, temperatures

must be high enough to induce melting. One major problem associated with understanding 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.

Par tial 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 is

required 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/or

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

S eg r e g a t i o n a n d R i s e o f M a g m a s T h r o u g h T h e M a n t l ee gr e ga t i on a n d R i s e o f M a g m a s T h r o u g h T h e M a n t l eSeveral mechanisms of magma rise through the mantle have been visualized. These

processes include:

(a)Deep Segregation: The melt forms along a dendritic network of joints 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 of

earlier bodies.(b)Diapiric Rise: A density reversal can lead to what is known as Rayleigh-

Taylor instability in which lighter underlying material first collects in 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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R i s e o f M a gm a s T h r o u g h B r i t t l e L i t h o sp h e r ei se o f M a g m a s T h r o u g h B r i t t l e L i t h o sp h e r eIt is difficult to determine the level at which the lithosphere deforms by brittle fracture rather

than by plastic flow - a depth represented by earthquake foci. There is strong evidence, in the 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 and

horizontal 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 roof

rocks. 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 heated

groundwater propagated downward as pressure is released onprogressively deeper gas-charged horizons.

F l o w o f M a g m al o w o f M a g m aKnowing the rheological or fluid properties of magmas, we might be able to apply basic

fluid dynamic principal to predict flow regimes of intrusive and extrusive magmas under 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 = (r4)/8L

where Q is the volume flow rate in cm3 sec-1, P is the pressure drop in bars, r is the channel radiusin cm, 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 filled

with basaltic magma ( = 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/secor 3.76 km3/day.

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This simple calculation is important in that it illustrates that movement of large quantities of magmain short periods of time is entirely feasible.

Natu r e of Flow Regime

The type of flow imposed on a magma, that is, laminar or turbulent flow, is also of

interest. For example, in the case of an initially heterogeneous magma, the liquid would 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 = (2Q)/r

where is the density of the fluid. Turbulent flow occurs when Re > 2000. For the previousexample, with = 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, then

the 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 Insta bilities

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 convection

accelerates 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 Tg)/K

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where L is the height of the layer in cm, T is the coefficient of thermal expansion, g is theconstant of gravitational acceleration (980 cm/sec), is the vertical temperature gradient in K cm-1, is the kinematic viscosity (), K is the thermal conductivity of the magma in cal gm-1 K -1, and 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, , 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.

H i g h - L e ve l R e s er v o ir s a n d S u b v o lc a n i c S t o c k si gh - L e v el R e s e r v o i r s a n d S u b v o lc a n i c S t o c k sThe erosion of extinct volcanoes reveals the presence of simple and multiple stocks of

medium- to coarse-grained rocks. Generally, the stocks are 1- to 10-km-wide, circular to oval 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 magma

bodies 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 explosive

eruption 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 ch a n i sm sI I . E r u p t i ve M e ch a n i sm sO p e n i n g O f V en t sp e n i n g O f V e n t s

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.

M e c h a n i s m s of E x p l os i ve E r u p t i o n se c h a n i s m s o f E x p l os i ve E r u p t i o n sAll explosive eruptions involve the sudden release of energy by gas under pressure, but the

way gas expansion acts on magmas varies widely. The explosivity of a volcanic eruption does 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.

N a t u r e o f t h e G a s eo u s E r u p t i ve C o lu m na t u r e of t h e G a s e ou s E r u p t i ve C o lu m nTo understand fully eruption mechanisms it is useful to examine the characteristics of the

eruption column and how it varies as magma reaches to the vent:

(a) Temperature Relations: Exsolution and expansion of gas significantly 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 300Ccooling 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 100C. 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 magma

approaches 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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B u b b l e N u c le a t i o n A n d G r o wt hu b b l e Nu c l ea t i on A n d G r o wt hIn order to understanding the mechanisms of explosive eruptions, it is useful to consider

the manner in which gas exsolves from the magmatic liquid. Even in the most viscous magmas, therate of bubble nucleation is very high. In order to evolve and grow, gas bubbles must reach an

initial size that balances the surface tension () of the magma at the gas-liquid interface.The pressure of gas inside the bubble acts over a cross-section r2, and is balanced by

surface tension around the circumference of its walls in the same cross-section (2r). Therefore,the gas pressure must exceed a value of 2/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. Exsolution

and expansion of dissolved gases ultimately leads to disruption of the coherent magmatic liquid.

Pressure Relations

The principal factor controlling the violence of explosive eruptions is the magnitude 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:(gh)

(b) the pressure required to drive the magma through the conduit:

P = 12Vh/r3 in cylindrical conduitsP = 12Vh/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 exerts

a pressure over the cross-sectional area of the bubble, so that the totalpressure from surface tension through the vesiculated liquid is:

P = 2n2/3/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 > 2n2/3/r +

where 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 = 4/ 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 on

magma 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.

E j e c t i on O f P y r o cl a s t ic M a t e r i a lj e c t i o n O f P y r o c la s t i c M a t e r i a lAs mentioned previously, the ability of the eruption column to carry in suspension and

eject fragments of disrupted magma is determined by the column density. The nature of ejecta 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.

Ej ection 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 value

when 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 and

thermal 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.

Er uption 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 sV . L a v a F l o w sLava flows are the products of extrusion of a coherent magma body onto the Earth's

surface. The external forms and internal structures of lava flows are the result of both the physical

properties of the magma and the external environment in which extrusion takes place. The 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.

V o l u m eo l u m eBasalts 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.

L e n g t h a n d T h i c k n e sse n g t h a n d T h i c k n e ss

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)BasalticLavas: 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 flow

has 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. Hornblende

andesite lavas tend to form short stubby flow that have the form of domes.(c)Dacitic and Rhyolitic Lavas: Siliceous lavas are short and thick. Few of these

lavas travel more than 1 or 2 km, and many come to rest on 30-40 slopes.Some of the largest known siliceous lavas include:

(a) The Big Obsidian flow (Medicine Lake) - 1 km long,(b) Glass Mountain Obsidian Flow - > 5 km in length, and,(c) Ring Creek dacite - 27 km long and up to 250 km thick.

V e l o c i t y o f F l o we l o c i t y o f F l o wThe 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) the

slope 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 Lavas

usually on the order of 10's or 100's of meters/hour.

D is ch a r g e R a t e si sc h a r g e R a t e sThe discharge rate of lava flows from the volcanic vent depends principally upon the

fluidity of the magma, and size and dimensions of the conduit. Like flow velocity, discharge 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 Parcutin - 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 ic a l P r o p e r t i es o f L a v a sh y s ic a l P r o p e r t i es o f L a v a sTempera tur e 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 than1100C.

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 of

cooling 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/())0.5,

where Q is the heat flux per unit time t, K is the conductivity of the ground, 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 500-650dark red 650-800bright red 800-1000orange 1000-1150yellowish-white 1150-1300

Viscosity

There are very few measurements of the viscosity of flowing lavas, but this property may

be estimated from the relation:

= gsinAd2/3V

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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 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: = 2.611.5

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 gy O f L a v a F l o w so r p h o l og y O f L a v a F l o w sLava flows exhibit a variety of morphologies that depend on the magma viscosity and the

external environment. Several types of lava flows are recognized to occur in lavas of different 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 lava

crust (rootless vents). They are formed by the discharge (locally explosively)of clots of lava that adhere to earlier clots to form a pile of welded ragged-surface fragments in a deposit called agglutinate.

(f) Lava tree molds and casts: These are molds formed when lava flows 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 passes

across 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 central

gutter.(c)Lava Lobes: These features represent lava tongues that have generally

developed along flow margins after the levee is breached.

(d)Accretionary Lava Balls: These structures form, like snowballs, by the 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 La va

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 cubes

in 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 former

have several distinguishing characteristics:(a) Pillows are rimmed by chilled glass selvedges, formed by rapid cooling of

lava by the surrounding water.(b) Pillows vary from a few cm to several m in diameter, and are generally

spheroidal, ellipsoidal, or may be flattened in cross-section.(c) Pillow tops are usually convex upward, whereas their undersurfaces may be

flat, concave upward or project downward between the underlying pillows.(d) Where gas cavities are present, these structures tend to be located within the

upper part of the pillows.

Pillow lavas are locally found in association with several other types of subaqueous 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 s. V o l c a n i c D o m e sOver or near the vent, extremely viscous lava tends to pile up into steep-sided heaps of

molten rock, known as domes or tholoids. Some domes also result from the bodily upheaval of

material filling the upper part of the conduit. This semi-solid to solid material is pushed up like 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 less

circular, 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 Vo lc a n i c D om e sx t e r n a l F e a t u r e s o f V o lc a n i c D om e sThe exteriors of volcanic domes are distinguished by several

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