Geo.416 Volcanology

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

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

I. Physical Nature Of Magmas . . . 4

Structural State of Silicate Melts . . . 4

Viscosity . . . 5 Controls on Viscosity . . . 5 Silica composition . . . 6 Temperature . . . 6 Time . . . 6 Volatiles . . . 7 Pressure . . . 7 Crystal content . . . 7 Bubble Content . . . 8 Yield Strength . . . 8 Specific Heat . . . 8 Thermal Conductivity . . . 8 Density . . . 9 Electrical Conductivity . . . 9

Seismic Wave Velocities . . . 9

II. Generation, Rise And Storage Of Magma . . . 10

Nature of Crust and Upper Mantle . . . 10

Heat Sources . . . 11

Mechanisms of Melting . . . 11

Partial Melting . . . 11

Segregation and Rise of Magmas Through The Mantle . . . 12

Rise of Magmas Through Brittle Lithosphere . . . 13

Flow of Magma . . . 14

Flow Rates . . . 14

Nature of Flow Regime . . . 15

Flow Instabilities . . . 15

High-Level Reservoirs and Subvolcanic Stocks . . . 16

III. Eruptive Mechanisms . . . 18

Opening Of Vents . . . 18

Mechanisms of Explosive Eruptions . . . 18

Nature of the Gaseous Eruptive Column . . . 19

Bubble Nucleation And Growth . . . 20

Pressure Relations . . . 20

Ejection Of Pyroclastic Material . . . 21

Ejection Velocities . . . 21

Eruption Energy . . . 22

V. Lava Flows . . . 23

Volume . . . 23


Velocity of Flow . . . 24

Discharge Rates . . . 24

Physical Properties of Lavas . . . 25

Temperature and Cooling of Lavas . . . 25

Viscosity . . . 25

Morphology Of Lava Flows . . . 26

Pahoehoe Lavas . . . 26

External Structures of Pahoehoe Lavas . . . 26

Internal Structures of Pahoehoe Lavas . . . 27

Aa Lavas . . . 28

External Structures of Aa Lavas . . . 28

Block Lava . . . 29

Internal Structures of Blocky Lavas . . . 29

Pillow Lava . . . 30

VI. Volcanic Domes . . . 31

External Features of Volcanic Domes . . . 31

Internal Structures of Volcanic Domes . . . 32

VII. Products Of Volcanic Explosions . . . 33

Terminology and Classification . . . 33

Origin . . . 33

Fragment Size . . . 33

Airfall Ash Deposits . . . 35

Dispersal . . . 36

Structures Of Airfall Deposits . . . 36

Morphology of Ash Particles . . . 37

Pyroclastic Flow And Surge Deposits . . . 37

Relationship to Topography . . . 38

Flow Units and Cooling Units . . . 39

Components . . . 40

Characteristics of Ash-Flow Deposits . . . 40

Internal Layering . . . 40

Gas-Escape Structures . . . 41

Textural Relationships . . . 41

Segregation of Crystals and Lithics . . . 42

Temperature Effects . . . 42

Welding and Compaction . . . 42

Structures Related to Temperature and Viscosity . . . 43

Classification and Nomenclature of Pyroclastic Flows . . . 43

VII. Laharic Deposits . . . 47

General Features . . . 47

Surface of Lahars . . . 47

Basal Contact of Lahars . . . 48

Components of Lahars . . . 48

Grain-Size Distribution . . . 48

Grading . . . 48

Fabric . . . 49


Internal Structure . . . 51 Maar Volcanoes . . . 51 Littoral Cones . . . 52 Shield Volcanoes . . . 52 Icelandic Shields . . . 52 Hawaiian Shields . . . 53 Galapagos Shields . . . 53 Composite Cones . . . 54 External Form . . . 54 Internal structure . . . 54 Growth Sequences . . . 55

Parasitic (adventive) Cones . . . 55

IX. Craters, Calderas, and Grabens . . . 56

Explosion Craters . . . 56 Collapse Craters . . . 56 Calderas . . . 56 Classification of Calderas . . . 57 Krakatoan Type . . . 57 Katmai Type . . . 58 Valles Type . . . 58 Hawaiian Type . . . 58 Galapagos Type . . . 59 Masaya Type . . . 59 Atitlán Type . . . 59 Cauldrons . . . 60 Volcano-Tectonic Depressions . . . 60 Resurgent Calderas . . . 60

X. Classification Of Volcanic Eruptions . . . 61

Nature of Vent . . . 61

Styles of Eruptive Activity . . . 61

Hawaiian Eruptions . . . 61 Strombolian Eruptions . . . 62 Peléean Eruptions . . . 62 Plinian Eruptions . . . 62 Vulcanian Eruptions . . . 63 Surtseyan Eruptions . . . 63 Appendices . . . 65

A. Pyroclastic Fall Deposits . . . 65

B. Pyroclastic Flow Deposits . . . 66

C. Pyroclastic Flow Deposit Characteristics . . . 67

D. Pyroclastic Surge Deposits . . . 68


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

G 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 s

I . 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, solidifies as a crystalline or glassy igneous rock. It is usually rich in silica and capable of flowing under moderate differential stress. Magmas may carry rock fragments or crystals in suspension, and they normally contain gaseous (volatile) components in solution.

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

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

S t r u c t u r a l S t a t e o f S i l i c a t e M e l t s

S t r u c t u r a l S t a t e o f S i l i c a t e M e l t s

Modern concepts of silicate liquid structure are based on the Zachariasen Model. The atoms are bonded by forces similar to those between atoms of crystals, but lack long range periodicity and symmetry. The magmas have silica (and alumina) tetrahedra linked (or polymerized) in three-dimensional networks in which (bridging) oxygen atoms are shared by two or more tetrahedra; the Si and Al cations are termed "framework cations." Other cations enter the melt in limited amounts as independent ions occupying positions between tetrahedra, and modify the basic structural framework and its physical properties; these cations, Ca, Mg, Fe, Mn, Ti, Na, and K, are termed "framework-modifying cations."

The framework-modifying cations can be accommodated in amounts of up to about 20 cation percent before the basic framework breaks down into smaller geometric units. In breaking liquid continuity into smaller units, the framework changes from an extensive network of tetrahedra, all of which are linked by shared O atoms to smaller units with lower Si:O ratios until, when more than 66% of the cations are framework modifiers, the liquid consists of separate tetrahedra not directly linked to each other.

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


V i s c o s i t y

V i s c o s i t y

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

s = so + η dm dt



where s is the total shear stress applied parallel to the direction of deformation; so is the yield strength of the fluid or the stress required to initiate flow; η is the viscosity, expressed in units called poises (dyne sec/cm2); dm/dt is the gradient of velocity dx/dt or strain rate over a distance Z

normal to the direction of shear; and, n is an exponent which has a value of 1.0 or less depending on the form of the velocity gradient.

For many fluids, this expression describes a linear relation between the strain rate (dx/dt) and shear stress parallel to the direction of shear. If a shear stress greater than the yield strength (s > 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 and returns to its original form after the stress is removed. Some fluids do not require application of some initial force before they are permanently deformed by shear stress parallel to the direction of shear. Such fluids are said to exhibit Newtonian behavior when n equals 1.0 and soequals zero.

Highly polymerized or non-Newtonian fluids (known as Bingham liquids) have a finite yield strength that must be exceeded before they can be deformed permanently. In other words, Bingham fluids behave elastically until their yield strength is exceeded.

Cooling and crystallizing magmatic liquids behave as newtonian fluids only until they contain approximately 20% crystals. Liquids with suspended solid particles may have a non-linear relation of shear stress to strain rate, for which the value of n is less than 1.0.

Controls on Viscosity

Various factors control magmatic liquid viscosity: composition (especially Si and volatiles), temperature, time and pressure, each of which effect the melt structure. Actually, the viscous behavior of complex silicate liquids, such as magmas, is difficult to predict, because no comprehensive theory explains the effects of major cations or temperatures of magmatic conditions.

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


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

Silica composition

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

Na:Si:O η (poises)

0:1:2 1010

1:1:2.5 28

2:1:3 1.5

4:1:4 0.2

The decrease in viscosity can be attributed to a reduction in the proportion of framework silica tetrahedral, and therefore, strong Si-O bonds in the magma.


Temperatures of erupting magmas normally fall between 700° and 1200°C; lower values, observed in partly crystallized lavas, probably correspond to the limiting conditions under which magmas flow. Low temperatures characterize silica-rich rhyolite magmas, whereas the highest temperatures are observed in basalts. Magmas do not crystallize instantaneously, but over an interval of temperature. Few magmas, however, have a wide enough range of crystallization to remain mobile at temperatures far below those at which they begin to crystallize or much hotter than those temperatures.

Temperature has a strong influence on viscosity: as temperature increases, viscosity decreases, an effect particularly evident in the behavior of lava flows. As lavas flow away from their source or vent, they lose heat by radiation and conduction, so that their viscosity steadily increases. For example:

a) measured viscosity of a Mauna Loa flow increased 2-fold over a 12-mile-distance from vent;

b) measured viscosity of a small flow from Mt. Etna increased 375-fold in a distance of about 1500 feet.

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


At temperatures below the beginning of crystallization, viscosity also increases with time. If magma is undisturbed at a constant temperature, its viscosity may continue to increase for many hours before it reaches a steady value. The viscosity increases with time results partly an increasing proportion of crystals (which raise the effective magma viscosity by their interference in melt flow), and partly from increasing ordering and polymerizing (linking) of the framework tetrahedra.



The solubility of gases in magmas varies with pressure, temperature and composition of both the gas and the magmatic liquid. Because the volume of a melt with dissolved gas is less than that of a melt and separate gas (vapor) phase, solubility increases as gas pressure increases. At constant gas pressure less than total pressure, any increased load pressure on the melt lowers solubility, because the volume of the melt with dissolved gas is greater than that of melt alone.

Vapor pressure increases with temperature, so that solubility of any volatile component generally decreases with temperature, except possibly at high pressure. Consequently it is difficult to predict how volatile content of magma varies with depth. Nevertheless, it has been shown that at constant temperature, solubilities of water in magmas with different compositions are not significantly different.

Nearly all magmas can contain more water or gases at depth than they can continue to hold in solution when they reach the surface. Basalts, however, normally contain less water than rhyolites simply because their temperatures are higher, and thus, as noted, lower gas solubility. Only limited data exists concerning the effect of volatiles (in particular F, Cl, S, H2S, SO2, CO,

and CO2) on magma viscosity. No doubt, the effect of dissolved water is to lower viscosity, the

effect being greater for silica-rich than silica-poor magmas:

Magma T (°C) 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 such polymerization is already minor or absent, there is little effect. F and Cl are though to considerably reduce magma viscosities; in contrast, CO2 increases polymerization, and therefore viscosity, in

melts by forming CO3-2 complexes.


The effect of pressure is relatively unknown, but viscosity appears to decrease with increasing pressure at least at temperatures above the liquidus. As pressure increases at constant temperature, the rate at which viscosity decreases is less in basaltic magma than that in andesitic magma. The viscosity decrease may be related to a change in the coordination number of Al from 4 to 6 in the melt, thereby reducing the number of framework-forming tetrahedra.

Crystal content

The effect of suspended crystals is to increase the effective or bulk viscosity of the magma. The effective viscosity can by estimated from the Einstein-Roscoe equation:


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-estimated value of 1.67.

Bubble Content

The effect of gas bubbles (vesicles) on the bulk viscosity of magmas can be variable, and depends on:

(1) the degree of bubble formation (that is, vesiculation); (2) the size and distribution of bubbles; and,

(3) the viscosity of the intervening melt.

Exsolution of water increases viscosity, but the exsolved vapor is a very low viscosity fluid; in basaltic magmas, the bubbles may enhance the already low temperature and composition controlled viscosity. Rhyolitic magmas have high viscosities irrespective of the degree of vesiculation, and only effect of high bubble content will be to reduce mechanical strength of the melt.

Y i e l d S t r e n g t h

Y i e l d S t r e n g t h

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

S p e c i f i c H e a t

S p e c i f i c H e a t

The specific heat (Cp) of magma, which is the heat required to change the temperature of the liquid 1 degree Celsius, is typically about 0.3 cal. gm-1. The specific heat contrasts greatly with

heat of fusion or crystallization, which is the heat that must be added to melt or removed to

crystallize a unit mass that is already at a temperature where liquid and solid coexist. Heats of fusion are typically about 65-100 cal. gm-1 at 1 atmosphere. Consequently, about the same amount

of heat is involved in crossing the crystallization interval, as in raising or lowering the temperature of the rock or liquid through 300°.

T h e r m a l C o n d u c t i v i t y

T h e r m a l C o n d u c t i v i t y

Igneous rocks and liquids are poor conductors of heat. Thermal conductivity depends on two heat transfer mechanisms:

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

The former declines and the latter increases as temperature increases and the melt structure expands. For rocks, the two effects balance each other up to their melting range. At high temperatures, the thermal conductivity of mafic rocks normally declines at an increasing rate up to 1200°C, above which, radiative heat transfer increases as does total thermal conductivity. More


D e n s i t y

D e n s i t y

Magma densities range from about 2.2 gm cm-3 for rhyolite to 2.8 gm cm-3 for basalts,

illustrating a close density-melt composition relationship, primarily reflecting the influence of higher concentrations of Fe, Mg and Ca cations in basalts. In contrast, magma density decreases with increasing temperature and gas content. These densities increase a few percent between liquid and crystalline states.

The temperature dependence of magma density is given by the coefficient of thermal expansion, about 2-3 x 10-5 deg-1 for all compositions. The pressure dependence of magma

density is given the compressibility or fractional volume change, ∆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


E l e c t r i c a l C o n d u c t i v i t y

E l e c t r i c a l C o n d u c t i v i t y

Electrical conductivity, which is low in pure silica melts, increases with increasing abundance of metallic cations, especially alkali elements, and increases abruptly in the melting range.

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

S 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, then

decrease abruptly to 2.5 km sec-1 at higher temperatures. Shear or S-wave velocities are about 2-3


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 a

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 a

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

N a t u r e o f C r u s t a n d U p p e r M a n t l e

N a t u r e o f C r u s t a n d U p p e r M a n t l e

Most of what is known concerning the Earth's interior comes from geophysical measurements, and concerns:

(a) seismic wave velocities; (b) temperature;

(c) density distributions; (d) heat flow; and,

(e) mechanical properties.

Seismic velocities increase with depth within the Earth, but show abrupt changes at several depths interpreted to represent discontinuities in the composition or structural state of minerals. The most notable discontinuities are:

(a) Mohorovicic discontinuity (MOHO); (b) Low Velocity Zone (LVZ); and (c) Core-Mantle boundary

The seismic velocities are closely related to the density 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 density increases to about 3.4 gm/cc at depths around 70 km, remains constant between 3.45 and 3.63 to the base of the Low Velocity Zone. Both pressure and temperature increase with depth. The temperature increase (6°/km) in the crust is consistent with an average heat flow of 1-2 x 10-6 cal.

cm-2 sec-1, with the highest values associated with young crust. If temperature gradients measured

in the crust are projected downward, they rapidly approach temperatures for beginning of melting in the mantle near the Low Velocity Zone. The transmission of shear or S seismic waves, however, suggests the absence of large amounts of liquid, so that the temperature gradients must diminish with depth.

H e a t S o u r c e s

H e a t S o u r c e s

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


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 low abundances in probable mantle rocks, too low to yield through their radioactive decay the heat necessary to generate magmas. Moreover, it can be shown that the melting process scavenges these elements, and thus, depletes even more their abundances in the source region.

Mechanisms of Melting

A variety of models have been invoked to explain the source of heat required to induce melting within the Earth:

(a) Stress Relief: Pressure on the source region is released during tensional or compressional deformation of the overlying rock column.

(b) Thermal Rise to Cusp in the Melting Curve: Intersection of pressure-temperature conditions with the source rock melting curve under conditions where lowest temperatures on the solidus coincide with phase change boundaries.

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

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

(e) Mechanical Energy Conversion To Heat: Force required to move one rock surface over another without grinding and deformation converted to heat, because of thrust faulting, subduction, a propagating crack or flaw in the Earth's lithosphere, shear or Tidal energy dissipated in the solid earth. (f) Compositional Change: The addition or subtraction of material changes the

rock composition to a new composition whose solidus lies at a temperature less than the ambient temperature.

Partial Melting

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

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

(a) temperature; (b) pressure;


(c) volatile content;

(d) mineral composition of the source rock; and, (e) amount or degree of melting.

Once gravitational instability sets in, the melt separates from the solid (denser) residuum. Depending upon where separation occurs, the magma may ascend through ductile (mantle) and/or brittle (crust) domains within the Earth. The manner in which magma rises differs between these two domains.

S e g 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 e

S e g 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 e

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

(a) Deep Segregation: The melt forms along a dendritic network of joints and fractures in the zone of melting, and feeds into a smaller number of layer tributaries eventually forming a larger channel at higher levels. With melting concentrated along grain boundaries, melt migration is caused by a thermal or pressure gradient or by capillary effects. This migration the presence of a critical proportion of melt before solid/liquid separation occurs. Two factors which could provide the driving force following initial separation are:

(i) pressure resulting from volumetric expansion on melting, and, (ii) the buoyancy of the liquid.

Once the liquid has separated, it is unlikely that it maintains a temperature much higher than its surroundings, as it is cooled by adiabatic expansion and conduction to the wall rocks. If the liquid rises slowly through rocks that are below their melting temperature, the magma would crystallize quickly. Thus, magmas can only ascend once the temperature of their wall rocks have been elevated, and successive batches of magma must tend to follow paths of earlier bodies.

(b) Diapiric Rise: A density reversal can lead to what is known as Rayleigh-Taylor instability in which lighter underlying material first collects in localized bulges under the heavier layer. The low density layer moves upward at an accelerated rate until it forms a steep sided plume or vertical density current. The rate of ascent , size, and spacing of plumes is a function of density differences, and the viscosity of the overlying rocks. Little or no separation of melt occurs in the zone of melting. Instead, the crystal-liquid mush rises and separation occurs at shallow levels. There again must be a delicate thermal balance between the diapir and its surroundings. Otherwise, it crystallizes.

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


R i s e 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 s p h e r e

R i s e 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 s p h e r e

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

(a) Dilational Rise: This proposed mechanism by which magma may rise involves: (i) entrance of melt in fractures, and rise due to gravitational buoyancy; (ii) The fracture becomes extended vertically and/or horizontally along a plane normal to the minimum stress; and, (iii) The fracture closes behind the magma as it passes and pressure on the wall falls below the confining pressure, rebounding due to viscoelastic deformation. Such a mechanism may explain the limited duration of basaltic fissure eruptions and the apparent arrival of discrete batches. Many instances, however, exist where acid or volatile magmas have apparently risen as pipe-like intrusions with little or no evidence of horizontal deformation.

The ability of a magma to rise through brittle lithosphere is usually explained in terms of depth and density contrast with the overlying rocks. If the pressure on the magma is equal to the lithostatic load of overlying rocks, the magma can rise to a level determined by the density contrast. At a depth of 50 km, the lithostatic pressure can exceed the pressure of a vertical magma column enough to segregate liquid and cause it to rise. If the heights to which magmas can rise is solely dependent on the depth to source and a density equilibrium, it would be expected that magmas with deep sources would erupt at higher elevations, and vice versa. This is obviously not the case as demonstrated by volcanoes of the Mexican volcanic belt.

More important limitations to magma rise are probably the heat content, and rates of ascent and cooling, which in turn, depend on the size of the magma body. Another important factor is the stress regime, which governs the form of the intrusive bodies. The three basic magma stress regimes are:

(a) least principal stress is horizontal (dikes); (b)least principal stress is vertical (sills); and,

(c) the stresses (vertical and horizontal) are equal (pipes; random dikes and sills).

At relatively high magmatic pressures or at shallow depths where vertical and horizontal stresses are low and about equal on the surrounding rocks, the magma conduits tend to be cylindrical. Thus, the form taken by a magma body may change drastically during its ascent. It is likely that near the surface, a cylindrical pipe is the most efficient form of conduit, because flow velocity increases and heat losses decrease as the horizontal section increases


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

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

An excellent example of non-dilational rise is illustrated by the formation of diatremes, steep-sided, more or less cylindrical or funnel-shaped breccia pipes formed by penetration of crust by moderate-temperature, gas-rich magma (kimberlite and carbonatite). Two mechanisms may be capable of boring through the Earth's crust and creating diatremes:

(i) Highly energized gases of deep-seated origin bore through the crust, opening channelways for the rapid ascent of magma; or, (ii) Explosive eruption is triggered by vaporization of heated

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

F l o w o f M a g m a

F l o w o f M a g m a

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

Flow Rates

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

Q = (ΡΠr4)/8ηL

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

in cm, η is the viscosity of the fluid in poises, and L is the length of the channel in cm. Applying this relationship to a large (about 200 km3) simple funnel-shaped magma chamber which is filled

with basaltic magma (η = 300 poises) via a 3-km-long, 200-m-wide, cylindrical feeder pipe at its base and a pressure drop through the pipe of 1000 bars (1 kb/3.3 km), we find:

Q = (3.14 x 1000 x 1016)/(8 x 300 x 3 x 105) = 4.36 x 1010 cm3/sec


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

Nature of Flow Regime

The type of flow imposed on a magma, that is, laminar or turbulent flow, is also of interest. For example, in the case of an initially heterogeneous magma, the liquid would become effectively homogenized by turbulence. The conditions that determine laminar or turbulent flow can be determined by calculating the dimensionless Reynolds number, Re, which in terms of average flow rate is given by:

Re = (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 higher viscosity of acid magmas, however, renders turbulent flow unlikely in these cases. Because the viscosity of magmas normally exceeds 103 poises and velocities are rarely greater than a few

cm/sec, flow is probably laminar under most geologic conditions.

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

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

In many cases, it is likely that a zone of static liquid will form a layer between the moving liquid and its solid boundary. Heat transferred from a cooling magma to surrounding wall rocks also affects its behavior in other ways.

Flow Instabilities

When heat losses from the top or sides of a magmatic body cause a density difference in the liquid large enough to produce gravitational instability, the liquid overturns and free convection accelerates the rate of heat transfer. The onset of convection in an infinite horizontal layer of viscous fluid having an upper and lower surface is given by the dimensionless ratio of buoyant to viscous forces known as the Rayleigh number, Ra:

Ra = (L4α


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 same value calculated for magmatic bodies of most common shapes. For a magma body of given size and viscosity, the principal variable is thermal gradient, β, a function of heat loss to the top or sides of the magma body. For Ra < 1000, transfer of heat is predominantly by conduction; steady convective heat transfer sets in at approximately Ra > 10000, and strong eddying motion is attained when Ra = 100000. Bodies with thickness or radius greater than 10 m are likely to convect if their heat losses are those that would be expected at shallow crustal depths (10-5 to 10-3 cal cm-2 sec-1).

Clearly, the larger the magma and the lower its viscosity, the more likely convection occurs, but quite small bodies having high heat flux values, should also be quite unstable.

H i g h - L e v e l R e s e r v o i r s a n d S u b v o l c a n i c S t o c k s

H i g h - L e v e l R e s e r v o i r s a n d S u b v o l c a n i c S t o c k s

The erosion of extinct volcanoes reveals the presence of simple and multiple stocks of medium- to coarse-grained rocks. Generally, the stocks are 1- to 10-km-wide, circular to oval in cross-section, and grade upward into a maze of inward dipping sills, steep radial dikes, and cone sheets. Most of these intrusive rocks have made room for themselves by stoping rather than forcible intrusion. There is good evidence that these intrusive bodies were volcanic reservoirs, because compositional features of erupted materials indicate that most magmas tended to reside and equilibrate in such shallow reservoirs prior to eruption. Other than what we see within deeply eroded volcanoes, however, little is known concerning the volcanic reservoirs beneath active volcanoes, except what is indicated by geophysics:

(1) Seismic methods: These methods have been used to detect large magma bodies at depth because of the inability of the Shear or S seismic waves to be transmitted through liquids. The distribution of earthquakes generated within or directly below a volcanic structure may delineate:

(a) the boundaries of intrusive bodies, and

(b)the possible movement of magma within the subvolcanic plumbing system.

For example, a three-dimensional distribution of earthquake foci surrounds an aseismic zone, which may represent one or more bodies of magma beneath Kilauea. Several types of earthquakes of volcanic origin are recognized according to the location of their foci and the nature of earthquake motion:

(a) A-type volcanic earthquakes: These earthquakes take place in and beneath volcanoes at places deeper than 1 km, generally in the range from 1 km to 20 km. They are generally less than 6 in magnitude. (b) B-type volcanic earthquakes: These earthquakes originate usually in


magnitudes are generally extremely small. The earthquake motions consist mainly of vibrations with periods in the range of 0.2 sec. to 1.0 sec.

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

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

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

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

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


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

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

O p e n i n g O f V e n t s

O p 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) the fractures through which magma reaches the surface represent planes of dilation propagated ahead of slowly rising magma; (ii) the appearance of lava is preceded by a mild release of steam or heated groundwater; and, (iii) eruption typically involves extrusion of magma that is relatively rich in gas. The strength, porosity and water content of near-surface rocks, shape and dimensions of the vent, and the physical properties of magma have a greater influence on the eruptive behavior than the depth of magma origin. Few explosive events are singular in nature, but rather represent an erratic succession of surges.

Magma does not reach the surface unless it is sufficiently heated to remain fluid and to penetrate the overlying barrier of cold rocks and groundwater. In order for these conditions to be met, it appears that a minimum conduit width and flow rate of magma within the feeder dikes is required. The final ascent of magma to the surface is neither sudden nor violent, but rather is a steady process that accelerates after the surface. The accelerated discharge may be due to:

(a) reduced resistance to flow;

(b) reduced density caused by expansion and vesiculation; (c) educed heat loss to surrounding rocks; and,

(d) increased temperature resulting from shear heating adjacent to dike walls.

The spacing and duration of eruptions seems controlled by the rates of stress accumulation in the lithosphere. Eruptions cease not because of a lack of magma, but due to a reduction in pressure.

M e c h a n i s m s o f E x p l o s i v e E r u p t i o n s

M e c h a n i s m s o f E x p l o s i v e E r u p t i o n s

All explosive eruptions involve the sudden release of energy by gas under pressure, but the way gas expansion acts on magmas varies widely. The explosivity of a volcanic eruption does not correlate directly with either volatile or silica content of the magma alone: the lowest is in those of olivine basalts, but highest in those of basanites and lamprophyres. The major factors which determine the explosivity are:

(a) the rate of gas expansion, and,

(b) the manner in which expansion occurs.

These factors, in turn, depend upon the viscosity of the magma, and the way in which they vesiculate. The degree of vesiculation and gas expansion may vary throughout an eruption.

Following a period or repose, initial eruptions usually therefore involve a gas-rich magma. Thereafter, the volatile content declines as gases escape to the atmosphere, and viscosity increases as more gas-poor magma is tapped. Low-density gas, either juvenile (magmatic) or meteoric (groundwater), concentrates in the upper parts of the plumbing system or reservoir by diffusing through a narrow boundary layer, through convective processes or by vesiculation and rise of


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

Expansion accelerates the rise of magma, so that the pressure of the overlying rock column is reduced at a faster rate, and eruption ensues. This process by which a reduction of lithostatic pressure allows an increase in exsolution of gas from the magma is known as "second boiling". Vesiculation could also be initiated by convective overturning of an density-stratified magma, or by injection of hotter magma (remember that, in both cases, a resulting temperature increase decreases gas solubility).

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

N a t u r e o f t h e G a s e o u s E r u p t i v e C o l u m n

N a t u r e o f t h e G a s e o u s E r u p t i v e C o l u m n

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

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

(b) Density Relations: The density of the eruptive column influences its capacity to carry fragments suspended in the gas stream. The smaller particles are subject to drag forces larger than their inertial forces, and thus, have lower terminal velocities so that they behave like gas particles. Particles less than 0.1 mm in diameter have so low terminal velocities compared to the velocity of the gas stream, that they contribute to the effective density and viscosity of the eruption column. A greater proportion of fine particles therefore enhances the ability of the eruption column to support large clasts or fragments.

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


B u b b l e N u c l e a t i o n A n d G r o w t h

B u b b l e N u c l e a t i o n A n d G r o w t h

In order to understanding the mechanisms of explosive eruptions, it is useful to consider the manner in which gas exsolves from the magmatic liquid. Even in the most viscous magmas, the rate of bubble nucleation is very high. In order to evolve and grow, gas bubbles must reach an initial size that balances the surface tension (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 can form if the gas pressure is greater than 6 bars (dry) or less (water-saturated).

Phenocrysts (large suspended crystals) accelerate vesiculation because bubbles that nucleate on the crystal surface require less volume to reach a given radius. The surface tension at a gas-liquid interface increases with falling temperature, but may be offset by dissolved water. The exsolution of water vapor increases surface tension to different degrees in different magmas, which may explain why bubbles tend to expand intact in some magmas but coalesce in others. Exsolution and expansion of dissolved gases ultimately leads to disruption of the coherent magmatic liquid.

Pressure Relations

The principal factor controlling the violence of explosive eruptions is the magnitude of residual gaseous phase, when the magma approaches the surface. There are four components of pressure in the vesiculating magma:

(a) the pressure of the overlying magma column: (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 is the conduit radius or w is the fissure width.

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


where n is the number of bubbles per unit volume and r is their average radius. The excess pressure of the gas phase, ∆P, exerts a force per unit of cross-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 than

50 percent, this excess pressure need only be a bars in order for fragmentation of the magma to occur.

(d) The pressure required for the bubble to expand against the viscous resistance

of the surrounding liquid:

P = 4η/ r (dr/dt),

where (dr/dt) is the expansion rate of the bubbles. This pressure, which varies between 10-2 bars and several hundred bars, is strongly dependent on

magma viscosity. In a fluid basaltic magma, a bubble with a 1 cm radius can grow radially at a rate of 0.5 mm/sec, more than enough to accommodate gas expansion at low pressure, but in viscous magmas, the expansion rate is two to three orders of magnitude slower and the pressure buildup is greater. The final sizes and gas pressures of bubbles are mainly a function of magma viscosity: the effect of increased viscosity during exsolution arrests expansion when the volumetric ratio of gas to liquid is between 3:1 and 5:1.

The first and second pressure components decrease as the magma rises and expands, whereas latter components are small. After the magma has vesiculated to the point that it behaves as a compressible fluid, i.e. the gas forms a continuous phase in which silicate liquid is carried in suspension, the second component, the dynamic pressure, becomes dominant.

E j e c t i o n O f P y r o c l a s t i c M a t e r i a l

E j e c t i o n O f P y r o c l a s t i c M a t e r i a l

As mentioned previously, the ability of the eruption column to carry in suspension and eject fragments of disrupted magma is determined by the column density. The nature of ejecta and the manner in which it is thrown out of the vent during eruption depends on their origin:

(a) primary material derived from the magma, or

(b) lithic fragments derived from conduit walls, with most plucked from the sides of the vent but some brought from deeper levels.

The principal difference in behavior of these fragments is that the primary magmatic fragments are part of the moving gas stream, whereas the accidental blocks are accelerated from rest.

Ejection Velocities

The muzzle velocities of ejecta depend on the size and settling velocity of fragments in the gas stream. The ejection velocity is the difference between the velocity of the gas stream and the velocity with which fragments would settle under static conditions. The minimum ejection velocity


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

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 more above the horizontal, and increase with depth to the focus of explosions. Velocities calculated with this expression are less than the actual ejection velocity, because as soon as a block leaves the effect of the gas stream, air resistance reduces its range, especially when it is small and has little momentum.

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

Below a few centimeters diameter, fragment movement is strongly retarded by wind and thermal currents. Estimated ejection velocities are on the order of 500-600 m/sec. Lower velocities are produced by the convective rise of warm air and gas. These currents, which are capable of carrying only fine dust, may reach great heights above the volcano, but velocities rarely exceed a few 10's of m/sec. The heights to which the eruption cloud rises therefore is related to:

(a) the vent radius; (b) the gas velocity;

(c) the gas content of the eruption; and,

(d) the efficiency with which thermal energy is converted to potential and kinetic energy during interaction with the atmosphere.

In general, large eruption clouds that reach high attitudes are produced by large eruptions of fine particulate material.

Eruption Energy

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

(a) heat energy contained in the solid and fluid products;

(b) heat and mechanical energy required to heat subsurface rocks and vaporize meteoric water;

(c) mechanical energy expended by magma and gas expansion; and, (d) work done against gravity during ascent of the magma.


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

I 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's surface. The external forms and internal structures of lava flows are the result of both the physical properties of the magma and the external environment in which extrusion takes place. The principal physical property that determines the nature of the lava flow is the magma viscosity, which is itself influenced by both the chemical composition of the magma and its temperature. The rate of magma supply to the flow is also important.

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

V o l u m e

V o l u m e

Basalts are not only the most abundant lavas, but they are also the most voluminous. Ultrabasic lavas are rare, and the abundance of andesitic, dacitic and rhyolitic lavas decreases as the magma viscosity increases with increasing silica and alkali content. The volumes of most historical lava flows are generally measured in the 10ths or 100ths of cubic kilometers. Some of the largest known lava flows include:

(a) 1669 Mt. Etna lava - ≈1 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 a

few 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 s s

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

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

(a) Basaltic Lavas: Fluid basaltic lava flows in Hawaii extend for more than 35 km with an average thickness of 5 meters. Some Icelandic basalts can be traced 80 km, whereas several Columbia River plateau basalts extended for mor than 100 km from their source vents. One Columbia River basalt flow has been traced over an area of 130 X 240 km, and has a thickness between 30 and 50 meters.

The length of lava flows is determined largely by the magma effusion rate. A high effusion rate, where lava spreads rapidly from the vent, usually results in a single flow unit. A low effusion rate, in contrast, results in lavas


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

(b) Andesitic Lavas: Andesitic flows generally have thicknesses of up to 30 meters, and are usually 5 to 15 km in length. Pyroxene andesite lavas typically are more extensive than those of hornblende andesite. Hornblende andesite lavas tend to form short stubby flow that have the form of domes. (c) Dacitic and Rhyolitic Lavas: Siliceous lavas are short and thick. Few of these

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

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

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

V 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 of effusion, (ii) magma viscosity, (iii) volume of magma extruded, (iv) magma density, and, (v) the slope and nature of the channel in which it flows. As expected, flow velocity diminishes with distance from the source. A pronounced velocity gradient exists within lava flows, extending from the middle toward the top, bottom and sides. Without a surface crust, the fastest movement occurs in the upper and middle parts of the flow, but once a crust forms, the fastest-moving part moves increasing downward into the lava. Some typical flow velocities are:

(a) Basaltic Lavas

30-60 km/hr Hawaii 8-75 km/hr Vesuvius (b) Siliceous Lavas

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

D i s c h a r g e R a t e s

D i s c h a r g e R a t e s

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

(a) 1947 Hekla - 75000 to 1250 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 of basalt, but there are notable exceptions:

(a) Sakurajima - 1666 m3/sec, and,


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

P 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 only rhyolitic obsidians are aphyric, or free of crystals. Because of their low thermal conductivity and high specific heat, most lavas are well insulated and cool slowly. Relatively little cooling takes place through most of the course of the flow, especially if the eruption temperature is greater than 1100°C.

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

Q = 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 to the 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 thermal conductivity, Ts is the surface temperature, and To is the initial temperature of the ground. Owing

to the low thermal conductivity and thermal diffusivity of soils and rocks, heat losses due to conduction are only a degree or two per hour.

Lava temperatures can be measured with (i) an optical pyrometer in which the color of incandescent lava is compared to that of a glowing filament; (iii) a sheathed thermocouple, or (iii) infrared techniques. A rough estimate of lava temperature (°C) may also be obtained from the color of the flowing magma:

brownish-red 500-650° dark red 650-800° bright red 800-1000° orange 1000-1150° yellowish-white 1150-1300° Viscosity

There are very few measurements of the viscosity of flowing lavas, but this property may be estimated from the relation:


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

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

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 of the lava accounts for flow fronts and levees ceasing to flow laterally even though slope angles may be great enough.

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

M 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 the external environment. Several types of lava flows are recognized to occur in lavas of different bulk composition.

Pahoehoe Lavas

These lavas are characterized by smooth, billowy, ropy or entrail-like crusts of quenched lava. Based on external form, various subtypes of pahoehoe lava can be distinguished:

(a) Massive: The lava crust is about 3 to 15 m thick, and smooth over large areas.

(b) Scaly: The lava consists of many small lobes or flow units that overlap like fish-scales. These units, sometimes called pahoehoe toes, may be 3-30 m in width and up to 30 cm thick.

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

(d) Slabby: These pahoehoe lavas are characterized by broken crusts, forming slabs a few meters across and a few cm thick.

External Structures of Pahoehoe Lavas

These lavas types, depending on magma viscosity, may show a variety of small-scale surficial structures that include:

(a) Lava Coil: These structures, which typify Shelly subtypes, consist of coiled, rope-like strips of magma crust, a few cm to about a m in diameter and 5-30 cm in height. The coils develop along shear zones between relatively stationary and adjacent blocks, being moved by undercurrents.

(b) Lava Blister: A mound of continuous lava crust, a few mm to a m in height and width, caused by the accumulation of gas beneath the lava surface. (c) Tumulus: This dome-shaped structure, resembling an ancient burial mound


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

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

(e) Hornitos, driblet- and spatter cones: These are small mounds or chimney-like spires that are built over eruptive vents or more commonly cracks in the lava crust (rootless vents). They are formed by the discharge (locally explosively) of clots of lava that adhere to earlier clots to form a pile of welded ragged-surface fragments in a deposit called agglutinate.

(f) Lava tree molds and casts: These are molds formed when lava flows around standing trees. After flow level subsides, a hollow cylindrical column is left by the carbonation of plant material.

Internal Structures of Pahoehoe Lavas

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

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

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

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

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


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

(c) Lava Tubes: These structures, which range from a few cm to 30 m or more in diameter and from a few km to 20 km in length, develop by the flow of molten magma within a confined interior channelway. In the upper parts of a lava flow, migration of magma eventually becomes restricted to these channelways. Fast-moving flows are characterized by relatively straight lava tubes, whereas slow-moving flows tend to contain meandering and branching tubes. If lava drains out of the tube before complete solidification, it leaves strandlines on the tube walls. In cross-section, lava tube walls are marked by concentric layers of congealed lava. Completely filled tubes show concentric bands of vesicles, platy joints parallel to the walls, and/or radiating joint columns. Lava stalactites and stalagmites may form by dripping of still-fluid lava from the tube ceiling; some may consist of sulfate minerals or opal. (d) Pipe vesicles and spiracles: These gas cavities are formed when lava passes

across wet ground, generating steam. The steam bubbles rise into the lava and form lines of vesicles or small tubes, usually less than 0.5 inches in diameter. If the upper end of the gas tubes are bent in the direction of lava movement, they are called pipe vesicles, and have been cited as possible indicators of flow direction. Where the steam bursts upward into the lava, it explosively creates an irregular, up to 10 m diameter, cylindrical opening called a spiracle. The spiracle generally terminates within the flow rather than extending through it, and may contain mud blown up from the underlying ground.

Aa Lavas

Aa lavas are characterized by surfaces that are a jumble of rough, clinkery and spinose, fragments, small chips to blocks measuring meters, and grade downward into massive lava. Based on external form, various subtypes of aa lava can be distinguished:

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

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

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

Aa lavas flow like a caterpillar tread, dumping talus over the snout and then overriding their own debris. Hence, they consist of a central massive part between fragmental top and bottom.

External Structures of Aa Lavas

Aa lavas types, depending on magma viscosity, may show a variety of large-scale surficial structures that include:




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