Ch 04 field relationships

Igneous Structures and FieldIgneous Structures and Field
RelationshipsRelationships
Thanks to John WinterThanks to John Winter

The style of volcanic eruption, and the resulting deposits, are determined by the
physical properties of the magma, in particular, the viscosity and volatile (gas)
content.
Viscosity (the resistance to flow) is determined by the composition and
temperature of the magma.
The strong Si-O and Al-O bonds in silicate melts can link together (or
polymerize) to form extensive networks.
Melts, at least near their melting point, have structures vaguely similar to the
minerals that would crystallize from them (although less ordered and less
polymerized).
Remember that each Si atom is bonded to four oxygens in silicate minerals. Each
oxygen can be bonded to another Si (called "bridging" oxygens, which create
-Si-O-Si-O polymers) or, more weakly, to some other cation ("non-bridging"
oxygens).
Properties of magma and eruptive stylesProperties of magma and eruptive styles

Because melts approximate the structures of the corresponding minerals,
basaltic melts tend to have more isolated Si-O with no bridging oxygens (as
in olivine) and partially bridged -Si-O-Si-O- chains (as in pyroxene), than
fully bridged three-dimensional -Si-O-Al-O- networks (as in plagioclase).
Because they have much more feldspar, rhyolitic melts tend to be
dominated by three-dimensional -Si-O-Al-O- networks.
The more polymerized -Si-O-Al-O- networks means that there are more
strong interconnected bonds in the rhyolitic melt, and thus higher viscosity.
Greater viscosity thus generally correlates with higher silica content.
Also, for a given SiO2 content, polymerization is weaker at high
temperatures and increases toward the melting point and with
increasing crystallinity.

When crystals form, the viscosity of resulting melt-crystal mixtures increases
rapidly as they cool, owing perhaps to the rigid crystals or surface adsorption
effects.
When stationary, such crystal-melt mixtures also develop a certain resistance to
induced flow. This resistance, called yield strength, must be overcome, before the
material can deform and behave as a viscous fluid again.
H2O and alkalis have the ability to break down some of the polymerized networks
into shorter segments, thereby reducing viscosity.
Volatiles are an important constituent of magmas for yet another reason.
As the magma rises, the pressure is reduced, and the volatile constituents escape
from solution in the magma and expand.
The gas pressures can become higher than the confining pressure of the
surrounding rocks at shallow levels. This can contribute to very explosive
eruptions.

H2O and CO2 are the dominant volatile species, but
there may also be SO2, H2, HCl, C12, F2, and a
number of other constituents.
Volatile content may range from less than 0.5 wt.
% in some basaltic magmas to over 5 wt.% in
some rhyolitic ones, and considerably more in
carbonatites and some other exotic melts.

Igneous Structures and FieldIgneous Structures and Field
Calculated viscosities of anhydrous silicate liquids at one atmosphere pressure, calculated by the method of Bottinga
and Weill (1972) by Hess (1989), Origin of Igneous Rocks. Harvard University Press. b. Variation in the viscosity of
basalt as it crystallizes (after Murase and McBirney, 1973), Geol. Soc. Amer. Bull., 84, 3563-3592. c. Variation in
the viscosity of rhyolite at 1000o
C with increasing H2
O content (after Shaw, 1965, Amer. J. Sci., 263, 120-153).

The combination of viscosity and volatile content determines whether a
volcanic eruption will be violent or quiescent.
Volatiles diffuse fairly easily through silicate magmas, especially if
polymerization is low.
Because volatiles are also of low density, they concentrate at or near the top
of shallow magma chambers.
As a result, the initial stages of most eruptions are more violent than the later
stages, because the volatile-rich upper portions are the first to be expelled.
Pressure keeps the volatiles dissolved in the magma while at greater depths, but
they are released when the magma approaches the surface and pressure is
relieved, in a manner similar to the uncapping of carbonated soda. If the
viscosity is low enough, the volatiles can escape easily.

The 1883 explosion of Krakatau in Indonesia
was heard near Mauritius, over 4800
kilometers away in the Indian Ocean, and
spread a cloud of ash around the globe,
lowering atmospheric temperatures
everywhere for several years.

Although most of the dissolved volatiles are expelled during pressure release
associated with the initial eruption, some remain in the magma.
Any dissolved gases remaining in the lava after the first eruption will leave solution
and form bubbles, called vesicles, which rise and concentrate toward the surface of
lava flows.
Highly vesicular basalt usually results from rapid vesiculation during explosive
eruptions; it is called scoria.
More viscous magmas, such as rhyolite can become so vesicular that blocks of the
cooled rock actually float on water.
This light-colored frothy glass is called pumice, and it generally forms as bits of
magma included in explosive ejecta, but also at the top of some rhyolitic flows.
Long after solidification of lava flows, vesicles may fill with later minerals
deposited by hydrothermal solutions.
Such filled vesicles are called amygdules and are common in mafic flows.

Basically two requirements must be satisfied if magma is to extrude,
either directly from its source in the deep crust or upper mantle where it
was generated or from a staging chamber in the shallower crust.

First, there must be an opening to the surface from
the buried magma body.

Second, magma must be able to move and be
propelled through the opening.
These are not necessarily independent of one another
and, in fact, are usually related.
Requirements of magma extrusionRequirements of magma extrusion

Several mechanisms, all of which depend on development of overpressure
in the subterranean magma body, allow venting of magma:
The first mechanism
 Independently of any exsolving and expanding volatiles, a buried mass of
magma may have the capacity to rise and even fracture the overlying
rocks by virtue of its buoyancy.
 Thus, in extensional tectonic regimes basaltic magma in upper mantle or
deep crustal reservoirs can invade subvertical fractures of its own making,
ascend to the surface, and extrude.
 This mechanism probably accounts for most extrusions of basaltic
magma.

Several mechanisms, all of which depend on development of
overpressure in the subterranean magma body, allow venting of
magma:
The second mechanism
After buoyant ascent from its source through dikes or as diapirs,
magma stored for a time in a shallow crustal chamber can
subsequently erupt once its volatile fluid pressure or buoyant force
exceeds the tensile strength of the roof rock overlying the chamber,
causing the roof to rupture and then allowing the gas-charged magma
to erupt.
This can happen in the following ways
As a stationary magma cools and crystallizes
feldspars, pyroxenes, and so on, the residual melt
becomes saturated in volatiles. The resulting
volatile fluid pressure or the buoyancy of the
bubbly magma can drive eruption.
The magma may rise to still shallower crustal levels,
causing more volatiles to exsolve and expand in the
decompressing system.
Exsolution and bubble growth may be retarded in
rapidly ascending, viscous magmas so that eventual
pressure release is greater and explosive eruption more
violent.
Instead of the magma’s rising to shallower levels to
cause decompression, a stationary magma system may be
unroofed.
Mafic magma may be injected into the base of a
chamber of cooler, less dense, more silicic magma.
 Transfer of heat to the intruded resident silicic magma
may create enough additional buoyancy to cause eruption.
But probably more significantly, cooling and
crystallization of the mafic magma cause volatile
saturation and exsolution.
Released volatiles float into the overlying silicic
magma, oversaturating it and increasing the volatile
pressure and magma buoyancy.
External water in the ground or in lakes or
ocean may come into contact with buried
magma, absorb heat, and expand
explosively, blowing off the shallow cover
over the magma body.

Structures and Field RelationshipsStructures and Field Relationships
Figure 4.2. Volcanic landforms associated with a central
vent (all at same scale).

Central volcanoes includes all those structures formed by volcanic
processes associated with a main central vent or magma chamber.
In vent eruptions, lava issues from a generally cylindrical, pipe-like
conduit through a subcircular surface hole (the vent).
There may be a bowl or funnel-shaped depression at the vent, called
a crater.
The rising magma follows weakened zones, fractures, or fracture
intersections in the brittle shallow crust. Magma solidifies in the vent
and plugs it, strengthening the old weakness and forcing later
eruptive phases to seek other conduits nearby.
Central VolcanoesCentral Volcanoes

Basalt is very fluid and flows rapidly, even on gentle slopes, to form what are
known as shield volcanoes.
Magmas with higher silica contents are more viscous and cannot flow as far from
central vents.
In addition, such magmas are more explosive and the resulting volcanoes, which
are composed of both lavas and fragmental material, are referred to as
composite.
Silica-rich magmas, such as rhyolitic ones, are so viscous that they do not
generally flow far from their vent; instead, they develop large blisters or domes
above the vent.
Explosive activity is so common with these silica rich magmas that the bulk of
eruptive material may be in the form of extensive fragmental deposits (tephra)

Shield volcanoes range in size from a few kilometers across to the
largest vent landforms.
The lavas that comprise a shield volcano are predominantly basalts.
Because of the relatively low viscosity, flows predominate, and cover
a large area, producing a landform with a convex upward profile and
low slope (generally less than 10° and commonly closer to 2 to 3°).
Over some "hot spots“ basalts issue forth in vast quantities, producing
the largest single landforms on Earth.
Although many flows in shield volcanoes come from the central vent,
there may also be marginal flank eruptions.

Composite cones are built of both lava flows and fragmental material
ejected from vents during periods of explosive activity.
These materials are interlayered and because of their markedly
different resistance to weathering, produce very stratified-looking
rocks.
Consequently the name strato volcano is also used.
They are composed predominantly of andesite.
Their eruptive histories are generally complex, having repeated
eruptions of both flows and pyroclastics, the layering of which give
the volcanic form its name.
These steep-sided cones are usually slightly concave-upward and
have slopes up to 36°.

a. Illustrative cross section of a
stratovolcano. After
Macdonald (1972), Volcanoes.
Prentice-Hall, Inc., Englewood
Cliffs, N. J., 1-150. b. Deeply
glaciated north wall of Mt.
Rainier, WA, a stratovolcano,
showing layers of pyroclastics
and lava flows. © John Winter

Schematic cross section of the Lassen Peak area. After
Williams (1932), Univ. of Cal. Publ. Geol. Sci. Bull., 21.

Cinder cone, Veyo,Cinder cone, Veyo,
UtahUtah

1982 eruption of St. Helens1982 eruption of St. Helens

Scoria cones result from the collection of airborne ash, lapilli, and blocks as they fall
around a central vent in association with weak explosive activity. They are typically less
than 200-300m high and 2 km in diameter, and generally last only a few years to tens of
years.
A maar is typically lower than a scoria cone, and has a much larger central crater, relative to the deposited
ring of debris. Maars result from the explosive interaction of hot magma with groundwater, which flashes
to steam. Such explosions are called hydromagmatic, or phreatic eruptions.

Tuff rings form when rising magma (usually
basaltic) comes closer to the surface than with a
maar before interacting explosively with shallow
ground or surface water. They also involve a
higher ratio of magma to H2O than maars, forming
a subdued ring of scoria and ash that has a low
rim, and layering of pyroclastic material that dips
inward and outward at about the same angle.
Tuff cones are smaller than tuff rings, with
steeper sides and smaller central craters. They
form where magma interacts with very shallow
surface water. They seem to result from less
violent and more prolonged eruptions than maars
or tuff rings. They look like scoria cones, but have
bedding that dips into the craters as well as
outward.

a. Maar, Hole-in-the-Ground, Oregon (upper
courtesy of USGS, lower my own). b. Tuff
ring, Diamond Head, Oahu, Hawaii (courtesy
of Michael Garcia). c. Scoria cone, Surtsey,
Iceland, 1996 (© courtesy Bob and Barbara
Decker).
b
c
a

Scoria cone, Ascension Island, South Atlantic

Domes form when largely degassed, viscous, silicic magma, such
as dacite or rhyolite (less commonly andesite), moves slowly and
relatively quietly to the surface.
High viscosities prevent lava flowing rapidly away from vent;
thus, as long as there are no violent explosions, lava accumulates
over the vent as blister or dome.
Domes vary considerably in diameter from few tens of meters to
several kilometers in diameter.
Lassen peak in California is 2km in diameter and 600m high.
DomesDomes

Domes may form early or late in an eruptive cycle, but they are typically
late.
After an early phase of explosive activity, the final (gas-depleted) magma
may inflate to a dome in the central crater.
The process is usually endogenous, in that the dome inflates by the injection
of magma from within.
Exogenous dome eruptions are events in which the later additions break
through the crust and flow outward.
The crust that forms on the dome brecciates as the dome inflates, giving the
surface a very rough texture. Brecciated blocks fall off and accumulate at the
base of the dome as an apron of talus.
DomesDomes

Schematic cross section through a lava dome.

Figure 4.8. Pressure ridges on the surface of Big Obsidian Flow, Newberry Volcano, OR. Flow direction is toward the left. © John
Winter and Prentice Hall.

A caldera is a large volcanic collapse
depression, more or less circular in form, the
diameter of which is many times greater than
that of any included vent. A crater may
resemble a caldera in form but is almost
invariably much smaller and differs
genetically in being a constructional form
rather than a product of destruction.
CalderasCalderas

Calderas form where a substantial volume of magma is
withdrawn from a subterranean magma chamber in a
geologically short time and the unsupported rock roof over the
evacuating chamber collapses into the growing void.
In more silicic cases, fractures reaching down to the magma
chamber reduce the pressure at the top of the chamber,
inducing rapid vesiculation, and a substantial proportion of the
magma may then escape along these fractures in the roof in the
form of pyroclastic activity.
CalderasCalderas

Development of the Crater Lake caldera.
After Bacon (1988). Crater Lake National
Park and Vicinity, Oregon. 1:62,500-scale
topographic map. U. S. Geol. Surv. Natl.
Park Series.

The largest calderas, some over 100km across, are associated with
tremendous pyroclastic eruptions of rhyolite.
The resulting calderas are too large to visualize at ground level, and
some have only been recognized recently from satellite imagery.
Examples of large calderas in the western United States are those at
Yellowstone, Valles Caldera in New Mexico, and Long Valley in
California.
Yellowstone contains several partially overlapping calderas. Such
occurrences are called caldera complexes.
There are 3 major eruptions at Yellowstone and collapses.
Following the collapse of the third caldera, the central part has risen
again (perhaps as new magma refilled a chamber below) to become
what is called a resurgent caldera.
CalderasCalderas

Figure 4.10. Location
of the exposed feeder
dikes (heavy lines) and
vents (V's) of the
southeastern portion of
the Columbia River
Basalts. Unshaded area
covered by CRB. After
Tolan et al. (1989), ©
Geol. Soc. Amer.
Special Paper, 239. pp.

Fissure eruptions occur as magma erupts to the surface along
either a single fracture or a set of fractures.
The planar conduits, when exposed by erosion, are filled
with solidified magma, and are referred to as feeder dikes.
Some fissure eruptions occur on the flanks of central vents,
where fractures typically accompany inflation of the edifice
as the magma chamber fills.
Fissure EruptionsFissure Eruptions
Such fractures may form singly, or in multiple sets (in a concentric or
radial pattern about the vent, or as parallel sets).
Fissure eruptions also occur in larger areas under-going regional
extension. Examples include the African Rift Valleys, the Basin and
Range, Iceland, and extensional basins behind volcanic arcs.

Lava flows are more quiescent than the dramatic explosive volcanic
eruptions, but they are the dominant form of volcanism on Earth.
Flows occur most typically in lavas with low viscosity and low
volatile content. They are thus most common in basalts.
Some flows, however, particularly those associated with flood
basalts, are of enormous size.
The lack of phenocryst alignment in many suggests that these flows
typically travel in a turbulent fashion, and come to rest before they
have cooled enough to be viscous and preserve any internal flow
features.
Types of basaltic lava flowsTypes of basaltic lava flows

Aerial extent of the N2 Grande Ronde flow unit (approximately 21 flows). After Tolan et al.

Basically, 3 types of basaltic flows can be recognized.
Three are typically subaerial, for which two have
names from the native Hawaiian tongue: aa and
pahoehoe .
The third flow type is subaqueous pillow lava. The
main subaerial aa and pahoehoe flows differ in surface
morphological characteristics and in the dynamics of
emplacement.
Types of basaltic lava flowsTypes of basaltic lava flows

In the early stages of a basaltic eruption, such as in Hawaii, magma
emerges as incandescent lava at about 1200°C.
This lava has a very low viscosity, and runs down slope in rivers with
initial velocities as high as 60 km/hr.
This runny lava cools and forms a smooth black surface, which may
develop a corrugated, or ropy, appearance.
Low-viscosity lava, especially basalt but also carbonatite, can produce
pahoehoe flows that consist of thin, glassy sheets, tongues, and lobes,
commonly overlapping one another.
Restrictions in downslope flow cause the glassy skin of the flow tongue to
wrinkle into ropelike festoons.
Downslope, lava pressure builds up within the rubbery skin, inflating the
sheet or tongue and causing breakouts of a new tongue.
Pa Hoe HoePa Hoe Hoe

It is common to see a lateral transition within a single lava flow from
pahoehoe morphology into a’a morphology.
The circumstances behind this transition are complex.
The field evidence may be summarized as follows:
l proximal pahoehoe lava often passes down-flow into distal a’a lava, though
in some cases a’a lava may be seen erupting directly from a vent;
l pahoehoe lava commonly changes laterally into a’a lava with increasing
distance from the vent, but a’a lava is never seen to revert to pahoehoe;
l pahoehoe morphology is seen only on low-viscosity basalt lavas, whereas
a’a surfaces are found on basalts and more evolved (more viscous) lava
types as well;
l pahoehoe lava often transforms into a’a where it begins to flow down a
steeper gradient.

Pahoehoe lava toes picked out by lighter - weathered margins, GranPahoehoe lava toes picked out by lighter - weathered margins, Gran
Canaria; hammer 28 cmCanaria; hammer 28 cm

A’a flow top on 2001 basalt lava, Mt Etna, SicilyA’a flow top on 2001 basalt lava, Mt Etna, Sicily

The hotter interiors beneath the skin of pahoehoe flows form an intricate
network of lava tube distributaries so the lava advances in multiple “fingers.”
Major feeding tubes in upstream parts of the flow can be several meters in
diameter and many kilometers long.
If, as commonly happens, the lava drains downslope from beneath the crusted
skin, the tube becomes an open cavity; these characterize exposed subvertical
sections through the interiors of solidified pahoehoe flows.

a. Ropy surface of a pahoehoe
flow, 1996 flows, Kalapana area,
Hawaii. © John Winter and
Prentice Hall.
b. Pahoehoe (left) and aa (right) meet
in the 1974 flows from Mauna Ulu,
Hawaii. © John Winter and Prentice
Hall.

Pa Hoe Hoe and aa flows.
A)Here the smooth ropy
surface of a still hot, steaming
pahoehoe flow covers an older
aa flow with typical rough
clinkery surface.
B) Detail of the pahoehoe flow
overlying the aa flow

Freshly extruded pahoehoe lava from May 1987 on Kilauea’s east rift zone, Hawaii. Various
scales of ropy structures indicate variable temperatures and rate of extrusion.
Note that lava that rose to fill the crack in the brittle crust on the right was cooler and more
viscous than the earlier lava and solidified before it could spill onto the surface of the flow.

Features that distinguish in cross-section between subaerial pahoehoe toes and
subaqueous pillow lava (a) Pahoehoe toes showing characteristic concentric vesicle
zones; black areas represent open tubes. (b) Pillow lavas showing characteristic radial
jointing, chilled glassy selvages (dashed lines indicate the inner margins), interstitial
sediment or hyaloclastite (grey), and voids within or between pillows (black).

Large blister formed on the surface of pahoehoe flow caused by pressure
of lava in the flow’s interior, Kamoamoa, Hawaii

As the lava cools further, and the viscosity increases, the flows begin to
move more slowly, and develop a thicker scoriaceous crust.
As the fluid interior continues to move, the crust breaks up into blocks of
clinkery scoria, which ride passively on the top.
Pieces also tumble down the advancing front.
The motion is like a conveyor belt, in which the surface slides beneath the
front of the advancing flow. Thus the blocks are found both at the top and
the base of the flow.
The rubble-like lava flows that result are called aa.
Block flows resemble aa but have a mantle of more regularly shaped
polyhedral chunks rather than jagged, highly vesicular, scoriaceous
clinker.
Block flows range from basaltic to highly silicic obsidian.
AaAa

Figure 4.12. c-e. Illustration of the development of
an inflated flow. In d, a thin flow spreads around a
rock wall. In (e), the flow is inflated by the addition
of more lava beneath the earlier crust. A old stone
wall anchors the crust, keeping it from lifting. The
wall can be seen in the low area in part (c). © John
ccThese flows begin as a thin pahoehoe
flow, perhaps as thin as 20 to 30 cm.
After the crust sets, continued lava is
added beneath it, and inflates the flow
internally to thicknesses as great as 18
m. As a result, the crust rises, cracks,
and tilts in complex patterns.

Pillow lavas are usually of low-viscosity basaltic magma formed
where it comes into contact with water or water-saturated
sediment, even in shallow intrusive situations (Walker, 1992).
Their most widespread occurrence is on the seafloor where they
have developed by extrusion along spreading ridges and on
seamounts.
Pillow basaltsPillow basalts
Although having the appearance in most
exposures, such as roadcuts, of a pile of
discrete, independent ellipsoids of pillow
shape and size, submarine pillow lavas in
some outcrops and especially those viewed
on the seafloor in three dimensions consist
of a tangled mass of elongate, grooved,
interconnected flow lobes that are circular
or elliptical in cross section

Subaqueous pillow basalts, Olympic Peninsula, Washington.
Hammer
for scale

Pillow lava on San Juan Islands

Llandwyn Pillow lava, North Wales, Europe

Pillow lavas, Maradihalli, 14km from Chitradurga

Pillow lava with associated hyaloclastite (Reykjanes peninsula, Iceland); hat (top left) 25 cm

Longitudinal view of pillow lava showing tube - like geometry, North Island, New Zealand; hammer
(lower right) 28 cm

Sub-aerial lava flows and some sheet like intrusions may develop a
characteristic jointing pattern called columnar joints.
As flood-basalts flows cool, they shrink and develop a characteristic
set of fractures which, because they break the rock into regular five- or
six-sided columns, are known as columnar joints.
Because the stress necessary to fracture the rock is generated by
cooling, the fractures propagate perpendicular to the cooling isotherms.
Cooling from upper surface is influenced by major fractures into which
water may percolate and cause additional cooling when the water is
converted to steam.
These “cold fingers” distort the isotherms, which cause the downward
propagating fractures to radiate out from these fingers.
Columnar JointsColumnar Joints

In the lower part of the flow, heat is lost by conduction into the
underlying rocks.
The cooling isotherms consequently parallel the lower surface
of the flow and fractures propagate up from the base as regular
columns.
Because cooling from the surface of a flow, especially if
augmented by evaporation of water, is faster than heat
conduction into underlying rocks, the downward and upward
propagating fractures meet at approximately one-third the height
of the flow.
Colonnade-entablature in Iceland

Columnar Joints, St. Mary Islands, Udupi

Columnar basalt in the Alcantara Gorges - Sicily - Italy

Colonnade and entablature in basalt lava, Isle of Staffa, Hebrides, off NW coast of
Scotland

Rosette jointing in basalt lava surrounding a tree cast, Ardmeanach, Isle of Mull,
Hebrides

The regular columns extending up from the base of the flow are
referred to as the colonnade and the upper zone of downward-
splaying fractures, as the entablature.
In the ideal case, there are 4 subdivisions:
1.A thin vesiculated and brecciated flow top
2.An upper
3.And lower colonnade with fairly regular straight columns
4.And central entablature that has more irregular columns that are
typically curved and skewed.
Various theories have been proposed for their mode of origin.

a. Schematic drawing of columnar joints in a basalt flow, showing the four common
subdivisions of a typical flow. The column widths in (a) are exaggerated about 4x. After Long
and Wood (1986).
b. Colonnade-entablature-colonnade in a basalt flow, Crooked River Gorge, OR. © John

The most widely accepted mechanism is by contraction of the
flow as it cools.
Because the top and bottom of the flow cool before the central
layer, these outer areas contract whereas the center doesn't.
This results in tensional stresses that create regular joint sets as
blocks pull away from one another to create polygons separated
by joints.
The joints propagate down from the top and up from the bottom
as cooling progresses toward the center, forming the column-like
structures in the colonnades.
However, formation of entablature is yet more difficult to
explain. Several theories exist.

The ideas generally relying on disturbed cooling surfaces or post-
joint deformation of the still ductile flow center.
Petrographic textural differences between the colonnade and the
entablature (Swanson, 1967; Long, 1978, Long and Wood, 1986)
suggest that the difference results from primary crystallization
effects.
The greater proportion of glassy mesostasis, and the feathery
Fe-Ti oxides in the entablature suggested to Long and Wood
(1986) that the entablature cooled more rapidly than the
colonnades, contrary to what one would expect for the more
insulated flow interior.

They proposed a model in which water infiltrated along the joints
(perhaps associated with lakes or floods) to cool the entablature
portion by convective water circulation through the upper joint
system.
Variation in the development of the subdivisions might then be
explained by the erratic nature of flooding, and the fanning of
columns in the entablature may be related to the preferential
percolation of water down large, early-formed joints.
Hence entablature undergoes rapid disequilibrium
crystallization whereas the colonnade crystallizes more
slowly under nearly equilibrium conditions.

Sometimes entablature forms at the top of flow.
In such cases Long and Wood (1986) have argued that the
fracturing and crystallization that characterize the
entablature are the result of quenching when water flows
across the surface of flow.
Some flows contain more than one level of entablature, which
they attribute to repeated flooding of the surface of the flow.
Lye (2000) has shown that these multitiered flows are usually
covered with fluvial sediments, which support the water-
quenching hypothesis.

The term volcaniclastic refers to any fragmental aggregate of
volcanic material, regardless of how it formed.
Autoclastic refers to volcanics that suffer a quiescent self-
imposed breakup, such as aa and block flows, dome talus
aprons, or gravitational collapse features.
Pyroclastic deposits are a subset of volcaniclastics that consist
of fragmented material formed from explosive volcanic activity
or aerial expulsion from a volcanic vent.
Other volcaniclastic deposits include volcanic mudflows, or
lahars, which form when volcanic debris mixes with sufficient
water, either rain or melted snow/ice, to mobilize.
Pyroclastic depositsPyroclastic deposits

Pyroclastic deposits may be deposited in water also. They may
then mix with the water to become lahars or water-laid deposits
(and are thus no longer pyroclastic by our definition).
Collectively, the pyroclastic particles that comprise the deposits
are called pyroclasts, and a collective term for the deposited
material is tephra.
Pyroclastic deposits are classified on the basis of mode of
transport and deposition.
They are divided into falls and flows.
Surges are distinguished by some geologists as a third
category of pyroclastic deposit.
Pyroclastic depositsPyroclastic deposits

They are produced by fallout from an eruptive jet or the plume of
an explosive eruption.
Plinian eruptions are combination of pyroclasts being forcefully
propelled upward during an eruption and then carried aloft by
hot convection.
The simplest falls accompany the small basaltic eruptions that
produce scoria cones and tuff rings/cones. The more dramatic and
larger eruptions produce a high ash plume that may reach heights
of 50 km.
As a result, fall deposits are well sorted, with grain size de-
creasing both vertically toward the top of a deposit and laterally
away from the source vent.
Pyroclastic FALL depositsPyroclastic FALL deposits

Ash cloud and deposits of the 1980 eruption
of Mt. St. Helens. a. Photo of Mt. St. Helens
vertical ash column, May 18, 1980 (courtesy
USGS). b. Vertical section of the ash cloud
showing temporal development during first
13 minutes. c. Map view of the ash deposit.
Thickness is in cm.

Approximate aerial extent and thickness of Mt. Mazama (Crater Lake) ash fall, erupted 6950 years ago. After
Young (1990), Unpubl. Ph. D. thesis, University of Lancaster. UK.

Maximum aerial extent of the Bishop ash fall deposit erupted at Long Valley 700,000 years
ago. After Miller et al. (1982) USGS Open-File Report 82-583.

They are left by dense ground-hugging clouds of gas-suspended
pyroclastic debris (mostly pumice and ash, with variable amounts of
lithic and crystal fragments).
Pyroclastic flows are controlled by topography, and the deposits
concentrate in valleys and depressions, unlike the uniform mantling
of ash falls.
Pyroclastic flows are hot (400-800°C). Velocities vary from 50 to
over 200 km/hr (the Mt. St. Helens blast was estimated at a max-
imum of 540 km/hr).
They can also travel long distances (typically a few kilometers, but
some deposits are found over 100km from source.
They are generated in several ways
Pyroclastic FLOW depositsPyroclastic FLOW deposits

Types of pyroclastic flow deposits.
a. Collapse of a vertical explosive or
plinian column that falls back to
earth, and continues to travel along
the ground surface.
b. Lateral blast, such as occurred at
Mt. St. Helens in 1980.
c. “Boiling-over” of a highly gas-
charged magma from a vent.
d. Gravitational collapse of a hot
dome.
e. Retrogressive collapse of an
earlier, unstably perched ignimbrite.

Field evidence clearly indicates that, when these flows come to rest
they are hot enough to cause some degree of welding of their
constituent glassy particles.
These welded ash-flow-tuffs are commonly referred to as
ignimbrites.
Some authors use the term block-and-ash deposits.
The rock name for a sample taken from the deposit is called a tuff.
Pyroclastic flow deposits at first glance appear to be unstratified and
poorly sorted, but closer examination generally reveals the presence
of stratification, cross bedding, internal erosion surfaces, graded
patterns, oriented particles and solid state deformation.
Pyroclastic FLOW depositsPyroclastic FLOW deposits

Section through a typical
ignimbrite, showing basal
surge deposit, middle flow,
and upper ash fall cover.
Tan blocks represent
pumice, and purple
represents denser lithic
fragments. After Sparks et
al. (1973) Geology, 1, 115-
118. Geol. Soc. America
Structures and FieldStructures and Field

These are formed from clouds of ash that move out
horizontally from a vent at hurricane velocities
following a volcanic explosion.
The passage of the cloud is short lived and thus only a
fraction of the total pyroclastic deposit associated with
an eruption is formed this way.
Bedding forms such as dunes and antidunes testify to
the high velocities of emplacement.
These clouds have high gas-to-solid ratio.
It was the switch from ash-fall to surge deposition
that resulted in most deaths of Pompeii in AD 79.
Pyroclastic SURGE depositsPyroclastic SURGE deposits

The generic term for an intrusive igneous body is a pluton, and
the rocks outside the pluton are called the country rocks.
The size and shape of plutons is generally somewhat speculative,
because erosion exposes only a small portion of most bodies.
The forms are grouped into tabular, or sheet-like, bodies and non-
tabular ones.
Further classification is based on specific shapes, and whether a
body cuts across the fabric (usually bedding) of the country rocks
or whether it follows the external structures.
Cross-cutting bodies are called discordant, and those that are
intruded parallel to the country rock structure are called
concordant.
Intrusive Processes and BodiesIntrusive Processes and Bodies

Tabular intrusive bodies are simply magma that has
filled a fracture.
A concordant tabular body is called a sill, and a
discordant one is called a dike.
A sill occurs when magma exploits the planar
weaknesses between sedimentary beds or other
foliations, and is injected along these zones.
A dike is a magma-filled fracture that cuts across
bedding or other country rock structure.
Dikes and sills are typically shallow and thin,
occurring where the rocks are sufficiently brittle to
fracture.
Tabular Intrusive BodiesTabular Intrusive Bodies

Figure 4.20. Schematic block diagram of some intrusive bodies.

Sill exposed in the Dry Valleys region of Antarctica. A large slab of sandstone can be
seen to have separated from the floor of the main sill and risen vertically, probably as a
result of buoyancy.

Although most dikes and sills are emplaced in a single
event, some may have a history of multiple injections.
More than one stage of injection may occur because
the dike or sill contracts as it cools, leaving a weakened
zone for later magmas.
Alternatively, the ductility contrast between a dike or
sill and the country rock might make the contacts
susceptible to localized deformation and later
magmatic injection.
A body is described as multiple, if all phases of
injection are of the same composition, and composite,
if more than one rock type is represented.

Dikes and sills can occur as solitary bodies, but dikes,
at least, more typically come in sets, reflecting the
tendency for fractures to form in sets as a brittle
response to imposed stresses over an area.
Genetically-related sets of numerous dikes or sills are
called swarms.
Dike swarms need not be parallel.
A radial dike swarm about a volcanic neck is illustrated
in the block diagram.
Such radial fractures are a common response to the
stresses imposed on the rocks above rising magma
bodies.

Figure 4.21. Kangâmiut dike swarm in the Søndre
Strømfjord region of SE Greenland. From Escher et al.
(1976), Geology of Greenland, © The Geological
Survey of Denmark and Greenland. 77-95.

a. Radial dike swarm around Spanish Peaks, Colorado. After Knopf
(1936), Geol. Soc. Amer. Bull., 47, 1727-1784. b. Eroded remnant of a
volcanic neck with radial dikes. Ship Rock, New Mexico. From John
Shelton © (1966) Geology Illustrated. W. H. Freeman. San Francisco.

A) The volcanic neck
towers 335m above the
plain. A resistant dyke
forms a prominent wall
extending south from
the neck.
B) The dyke is more
resistant than the
surrounding rocks and
consequently forms a
wall.
C) Satellite image showing
volcanic neck and radial
dykes. With of image is
10km

The formation of ring dikes and
cone sheets.
a. Cross section of a rising
pluton causing fracture and
stoping of roof blocks.
b. Cylindrical blocks drop into
less dense magma below,
resulting in ring dikes.
c. Hypothetical map view of a
ring dike with N-S striking
country rock strata as might
result from erosion to a level
approximating X-Y in (b).
d. Upward pressure of a pluton
lifts the roof as conical blocks
in this cross section. Magma
follows the fractures, producing
cone sheets. Original horizontal
bedding plane shows offsets in
the conical blocks

There are two principal types of concentric dikes: ring
dikes and cone sheets.
Ring dikes occur when the pressure exerted by the magma
is less than the weight of the overlying rocks. In this case,
cylindrical fractures will form.
If the overlying rock is slightly more dense than the
magma, cylinders of the roof will drop and the magma
will intrude along the opening fracture.
This is more likely to occur with less dense silicic
magmas.
Ring Dykes and Cone sheetsRing Dykes and Cone sheets

A fracture may also penetrate to the surface, in which case
a ring dike would feed a volcanic event, as in caldera
collapse.
The block that drops into the chamber may consist of
country rock, an earlier phase of the pluton itself, or
associated volcanics.
Ring dikes are either vertical or dip steeply away from the
central axis.

a. Map of ring dikes, Island
of Mull, Scotland. The classic
Tertiary ring dikes and cone
sheets. Note that there are two
centers, and several intrusive
phases with rings of different
compositions. Most of the
blocks within a ring consist of
earlier intrusives.

Cone sheets form when the pressure of the magma is
greater than the confining pressure of the overlying
rocks.
In this case, inwardly dipping fractures form, usually
as a nest of concentric cones.
Intrusion along these fractures produces a set of rings
at the surface which, in contrast to ring dikes, dip
inwards.
Cone sheets and ring dikes can occur together, where
they result from different phases of a single intrusion.

b. Cone sheets in the same area of Mull, after Ritchey (1961), British Regional Geology. Scotland, the Tertiary Volcanic
Districts. Note that the yellow felsite ring dike in part (a) is shown as the red ring in the NW of part (b). British Geological

The term vein refers to a small tabular body, whether or not it is
discordant or concordant.
It is typically used in association with ore bodies, where numerous
small offshoots from a pluton penetrate the adjacent country rocks.
These offshoots (veins) are typically quartz-rich and may contain
ore minerals.
The term is not recommended for other igneous bodies.
Rather, the terms dike or sill are preferable, no matter how small.
Dikes and sills can thus range in thickness from less than a
millimeter to over a kilometer, although most are in the 1 to 20m
range.
IntrusivesIntrusives

Figure 4.25. Types of tabular igneous bodies in bedded strata based on method of emplacement. a. Simple dilation (arrows) associated
with injection. b. No dilation associated with replacement or stoping. © John Winter and Prentice Hall.

A stock is a pluton with an exposed area less than 100 km2
, and a
batholith is a pluton with an exposed area larger than 100 km2
.
Distinguishing the two on the basis of exposed area is very simple
and useful in the field, but it is also unfortunate in that it is not
based on the size of the body itself, but in large part on the extent of
erosion.
It is possible that a small pluton that is deeply eroded would have an
exposed area greater than 100 km2
, and be a batholith, whereas a
much larger pluton, only barely exposed, would be a stock.
Also, several smaller stocks in an area may be so similar in age and
composition that they are believed to be connected at depth,
forming parts of a larger body. These separate extensions are called
cupolas.
Non Tabular Intrusive BodiesNon Tabular Intrusive Bodies

Block diagram several kilometers across, illustrating some relationships with the
country rock near the top of a barely exposed pluton in the epizone. The original upper
contact above the surface is approximated by the dashed line on the front plane. From
Lahee (1961), Field Geology. © McGraw Hill. New York.

Some types of stocks are genuinely smaller bodies, and
not simply limited exposures of larger batholiths.
Some stocks represent the cylindrical conduit and
magma chamber beneath volcanoes.
This type of stock is called a plug.
The exposed portion of a plug, commonly remaining
after the more easily eroded volcanics of the cone have
been removed, is called a volcanic neck.

In addition to the two "generic" types of plutons, stocks and
batholiths, there are a number of special types, based on shape rather
than size.
A laccolith is a concordant stock with a flat floor and an arched roof.
A lopolith is another concordant type of pluton intruded into a
structural basin. Both are essentially sills.
A laccolith is sufficiently viscous (and silicic) to limit magma flow
along the horizontal plane, and shallow enough to physically lift the
roof rocks.
Lopoliths are usually mafic, and characteristically much larger than
laccoliths. The Duluth gabbro complex, for example, is a lopolith over
300 km across.

Figure 4.26. Shapes of two concordant plutons. a. Laccolith with flat floor
and arched roof. b. Lopolith intruded into a structural basin. The scale is not
the same for these two plutons, a lopolith is generally much larger. © John

Gradational border zones between homogeneous igneous rock (light) and country rock (dark). After Compton (1962), Manual of
Field Geology. © R. Compton.

Figure 4.28. Marginal foliations developed within a pluton as a result of differential motion across the contact. From Lahee (1961),
Field Geology. © McGraw Hill. New York.

Figure 4.29. Continuity of foliation across an igneous contact for a pre- or syn-tectonic pluton.
From Compton (1962), Manual of Field Geology. © R. Compton.

Figure 4.31. a. General characteristics of plutons in the epizone, mesozone, and catazone.
From Buddington (1959), Geol. Soc. Amer. Bull., 70, 671-747.

Figure 4.32. Developmental sequence of
intrusions composing the Tuolumne
Intrusive Series (after Bateman and
Chappell, 1979), Geol. Soc. Amer. Bull.,
90, 465-482.
a. Original intrusion and solidification of
marginal quartz diorite.
b. Surge of magma followed by
solidification of Half Dome Granodiorite.
c. Second surge of magma followed by
solidification of porphyritic facies of Half
Dome Granodiorite.
d. Third surge of magma followed by
solidification of Cathedral Peak
Granodiorite and final emplacement of
Johnson Granite Porphry.
Structures andStructures and
FieldField

“inspite of much discussion, there is still no
consensus on how space is made for magmas in
the continental crust. A major problem involves 3
dimensional shape/geometry and in particular
evidence of the floors of plutons which are
curiously rarely observed”
Haerderle and Atherton (2002)
The process of magma rise andThe process of magma rise and
emplacement and “the room problem”emplacement and “the room problem”

Accounting for the “space” currently occupied by a large
intrusion is a key element in understanding the emplacement
mechanism: a large granite intrusion must have taken the
place of an equivalent volume of pre-existing crustal rocks,
so to where have those rocks been removed? If something
was “here” before that is no longer here, where is it now?
Though this “space problem” is not confined to granite
intrusions, the larger granite batholiths like the Sierra
Nevada of California and the Trans-Himalayan batholith
highlight it on the most spectacular scale.

One way to circumvent space problem is to
suppose that the granite was not introduced at all,
but has formed in situ.
19th
century French geologists developed a notion
called “granitization”, invoking “a process by
which solid rocks are transformed to rocks of
granitic character without passing through a
magmatic stage, a change mysteriously
accomplished by permeating fluids”
““Granitization”Granitization”

Migmatites, on the other hand, provide direct
evidence that granite can form more or less in situ.
A migmatite is an intimate outcrop scale mixture of 2
components: quartzofeldspathic layers, pods or veins
of granitic composition that cut or are interlayered with
darker gneisses.
They are found in regional metamorphic terrains that
have developed to sufficiently high grade to initiate the
partial melting – anatexis – of deep crustal rocks.

The lighter veins or layers (leucosome) are interpreted
as migrating partial melt frozen while still close to the
site of melting; the host gneiss (mesosome) may
develop darker selvages (melanosome) along the
contact with leucosome, which some interpret as the
complimentary solid residuum left behind by melt
extraction; the mesosome is regarded as source rock
that has yet to melt.

Intrusive igneous rocks are simply magmas that have not reached
the surface.
Plutonism lacks the drama of volcanism, because no one has ever
died from pluton emplacement, and there has never been a pluton
alert.
The volume of igneous rock in plutonic bodies, however, is
considerable, and there is a pluton of some sort beneath every
volcano.
In a sense we can think of these bodies as liquid that crystallized, or
"froze," in the plumbing system on its way toward the surface.
When magma forms by some melting process at depth, it segregates
from the residual unmelted solids to form a discrete liquid mass.

This mass is less dense than the surrounding solid, so it
becomes buoyant.
The buoyant magma body tends to rise, and, if the
surrounding material is sufficiently ductile, it is
generally considered to do so as a diapir.
A diapir is a mobile mass that rises and pierces the
layers above it as it does so.
Whenever a sufficiently ductile rock mass is covered
by denser rocks that are also ductile, diapirs are
capable of forming.
The process of magma rise and emplacement and “the room problem”The process of magma rise and emplacement and “the room problem”

• A well-documented example of diapirism is the formation of salt diapirs in areas
such as northern Germany, Iran, and the American Gulf Coast.
• In these cases, the salt beds (evaporite deposits) are covered by subsequent
sediments that compact and become denser than the salt beneath.
• Halite (salt) not only has a low density, but it will begin to behave in a ductile
fashion and flow under conditions of low confining pressure.

• After 5 to 10 km of overburden develops, the salt beds begin to flow, and the upper
surface develops irregular swells.
• Salt flows upward into these swells and they form diapirs that rise toward the
surface.
• If the diapir rises sufficiently above its source, and is not supplied by enough salt
flow from the evaporite bed, the column feeding the diaper stretches and constricts
until the body separates from the source bed.

• The inverted raindrop shape, with bulbous leading edge and
tapered tail, is characteristic of rising diapirs.
• Because buoyancy is the driving force that causes a diapir to
rise, once it reaches a level where the density of the surrounding
rock is the same as that of the salt, it will stop rising and spread
out laterally.

• Magma diapirs are believed to behave in a similar fashion:
Once the viscous drag between magma and wall rocks is
reduced sufficiently, buoyant magma could well up as a
viscous diaper resembling a salt dome, the up flow being
compensated by a peripheral down-flow of heated (and thus
ductile) country-rocks surrounding the diapir.
• Any bedding or other original fabric in the surrounding
medium is downwarped to form a peripheral synform owing to
removal of material from the source layer into the diaper.
• In the less ductile areas of the upper mantle and crust, rising
magma can no longer rise in diapiric fashion. Rather, it must
exploit fractures or other weaknesses in the rocks through
which it rises.

Open fractures and voids are limited to the very shallow near-
surface environment (a few tens of meters).
Below this, the rise of magma by simply filling such open voids is
not an option.
Rising magma can follow preexisting (closed) fractures, however,
by forcibly displacing the rocks that form the fracture walls and
following these planar conduits.
At depth, the ability of magma to force a fracture open is limited,
because the injecting pressure of the magma is seldom great enough
to significantly displace rigid rock walls that are forced together by
lithostatic pressure below a few kilometers depth.

Of course, if an area is undergoing regional
extension, the walls of any fractures will not be
under compression, and magma can force them
apart.
The number and width of dikes or sills in such
filled conduits depends upon the rate of extension,
and should be commensurate with typical plate
tectonic rates of less than ~ 3 cm/yr.

• The room problem becomes more troublesome for larger intrusive
bodies, which occupy a significant volume, and thus much more
rock must be displaced for them to move upward.
• Figure summarizes the proposed mechanisms by which a pluton
could make room and rise

• Plutons, such as laccoliths, can lift the roof (#1) by either folding or block elevation along
faults.
• Controversy exists over whether the lifting force of plutons is restricted to magma
buoyancy, which limits the ability of a pluton to lift the roof when it reaches the level at
which its density equals that of the country rocks.

• Lifting may be facilitated in such cases by magmatic overpressure, which might supply
additional depth-derived pressure.
• Roof doming to form laccoliths is limited to depths less than about 2-3km, where the
magmatic pressures can exceed the strength of the overburden (Corry, 1988). Similar
restrictions apply to block uplift along faults.

• Alternatively, magma may melt its way upward (assimilation: #2).
• The ability to melt the walls using magmatic heat is limited by the available heat of the magma.
• Thus the heat available for melting the country rocks does not exist in excess, and must be
supplied by the latent heat of crystallization of some portion of the magma, making it at least
partially solid, and therefore less mobile. This type of heat is necessarily limited.

• If the country rocks are sufficiently brittle, blocks of the roof over a rising pluton could
become dislodged, fall, and sink through the magma (#3).
•
This process is called stoping.
• Considerable evidence for stoping is to be found in the upper portions of many plutons, where
blocks of country rock are suspended in the crystallized igneous rock as rafts or xenoliths.

• Where the roof zone of a shallow granite or other felsic intrusion is
well exposed, one may sometimes see the roof rising or falling in
steps, and angular blocks of roof rock may be seen suspended in the
granite beneath.

• Stoping requires that the country rock be denser than the magma. Also, for stoping
to be effective, the stoped blocks must be large enough to sink rapidly enough in a
viscous magma, which can require blocks as large as tens of meters in granitic
melts.

• Cauldron subsidence and caldera formation are examples of large-scale stoping at shallow
depths.
• Stoping may be an emplacement mechanism at shallow depths, but cannot make the
original room for the pluton where there was none before.
• Ring and arcuate intrusions are examples.

• A combination of assimilation and stoping, called solution stoping or zone melting, may
operate at depths where the country rock is near the melting point.
• Such a process may be effective in the mantle and deep crust, where the magma rises as a
diapir.

• Ductile deformation and return downward flow (#4) are the mechanisms associated
with the rise of diapirs at greater depths, and are probably efficient where the viscosity
of the country rocks is low.
• The mechanism of roof yielding to this stretching would be ductile at depth, and more brittle
near the surface.

• Magma ascending through the crust becomes cooler and more viscous, and its upward
movement may therefore stall.
• Deeper fractions of the same magma stream continue to ascend and invade the cooler mass
above.
• This can lead to up-doming and radial inflation of the stalled magma body.

• The distinctive attribute of this process, known as ballooning, is that it causes a concentric
flattening foliation to develop, both in the outer zones of the pluton itself – as indicated by the
flattening of xenoliths and other strain markers – and in the surrounding country-rock
envelope, reflecting the outward directed compression imposed by the expanding core of the
magma body.(#5).

Sketches of diapirs in soft putty models created in a centrifuge by Ramberg (1970), In Newell, G., and N. Rast, (1970) (eds.),
Mechanism of Igneous Intrusion. Liverpool Geol. Soc., Geol. J. Spec. Issue no. 2.
• Here he modeled diapirs using soft materials of low density
covered by layers of higher density strata, and placed them in a
centrifuge to drive the diapiric motion.
• The low-density material rises as diapirs due to its buoyancy and
the ductility of all the material in the centrifuge

Figure 4.35. Sketches of diapirs in soft putty models created in a centrifuge by Ramberg (1970), In Newell, G., and N. Rast, (1970)
(eds.), Mechanism of Igneous Intrusion. Liverpool Geol. Soc., Geol. J. Spec. Issue no. 2.
• Upon reaching a level where the surroundings are less dense,
however, the buoyancy of the diapir is reduced or lost entirely, at
which point they tend to spread laterally and become more broad
and thin.

Figure 4.36. Diagrammatic cross section of the Boulder Batholith, Montana, prior to exposure. After
Hamilton and Myers (1967), The nature of batholiths. USGS Prof. Paper, 554-C, c1-c30.
• If this behavior is typical, even in the more brittle epizone,
batholiths may be much thinner than we have imagined, and the
amount of room required for their emplacement is greatly reduced.
• The above figure is a proposed cross section of the Boulder
Batholith in Montana, which suggests that it is less than 10 km
thick.
• Other plutons are now suspected of being this thin or even thinner.
• If this is true for most batholiths, it certainly alleviates the need for
great amounts of room, and lessens the room problem
considerably.

Figure 4.37. Possible methods by which a large batholith may grow by successive small increments over millions of years.
Magma rises initially as a series of dikes in an extensional terrane. Each dike spreads laterally as a thick sill upon reaching a
level at which it is no longer significantly buoyant. Room may be created by: a. lifting the roof rocks if the overburden is small,
b. depressing the chamber floor as magma is displaced upward and withdrawn from below (Cruden and McCaffrey, 2001;
Cruden, 2005), or c. some more irregular and sporadic process. Image courtesy of John Bartley.

Figure 4.38. Schematic section through a hydrothermal system developed above a magma chamber in a silicic volcanic terrane. After
Henley and Ellis (1983), Earth Sci. Rev., 19, 1-50. Oxygen isotopic studies have shown that most of the water flow (dark arrows) is
recirculated meteoric water. Juvenile magmatic water is typically of minor importance. Elsevier Science.

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