Wilson200722222222222

CHAPTER SEVEN
Active continental margins
7.1 Introduction
In Chapter 6 we considered the simplest type of
subduction-related magmatism produced as a con-
sequence of the subduction of one oceanic plate
beneath another. Now we shall focus our attention
on the more complex case in which the overriding
plate is a continental one. Magmas generated in this
tectonic environment occur along the west coast of
the Americas, Japan, Sumatra, Alaska, New Zea-
land and the Aegean (Fig. 7.1).
Since the early days of plate tectonics, the South
American Andes have been cited as the type
example of an ocean-continent collision zone, or
active continental margin (Mitchell & Reading
1969), and much ofthe discussion in the following
sections will be based on Andean data. A volcanic
arc developed upon an uplifted surface of Precam-
brian and Palaeozoic rocks along much of the
Pacific margin of the Americas by the late Triassic
or early Jurassic (Dalziel 1986) and volcanic activity
has been essentially continuous to the present day
along different segments of the plate margin.
However, compared to North America, the South
American continental margin has been a compara-
tively simple active margin since Triassic times and
consequently it may be utilized to develop general
models for a variety of geological processes which
may then be applied to more complex tectonic
situations.
The orogenic andesite association characteristic
of island arcs (Ch. 6) also typifies the volcanism of
active continental margins and, in many respects, is
broadly similar, although the passage of magmas
through thick continental crust produces added
complexities. Although it was once considered that
M. Wilson, Igneous Petrogenesis
© Springer 2007

192 ACTIVE CONTINENTAL MARGINS
ANTARCTIC PLATE
Active continental margins Continental micro-plates
1 Andes 3 Cascades 5 Japan 7 New Zealand
2 C e
ntral America 4 Alaska 6 Sumatra 8 Aegean
Figure 7.1 Location of the major active continental margins and subduction systems involving continental micro-plates.
such margins were dominated exclusively by calc-
alkaline rocks (Baker 1982), it is clear that the four
main magma series recognized in island arcs
(low-K, calc-alkaline, high-K and shoshonitic) are
all represented. Additionally, alkaline lavas are
often closely associated with the calc-alkaline volca-
nics, but generally form a separate zone of activity
to the landward side of the volcanic belt (Thorpe et
al. 1982). These may not necessarily be directly
related to the subduction system being formed in an
extensional regime similar to that of back-arc basins
(Ch. 8).
One of the most conspicuous differences between
the island-arc and continental-margin calc-alkaline
series is the greater abundance of more silica-rich
magmas (dacites and rhyolites) in the latter. Much
of this additional volume of acid rock occurs as
pyroclastic flow material (ignimbrite) and appears
to have a particular association with zones of
thickened continental crust. It is therefore a distinct
possibility that these acid magmas are derived, at
least in part, by partial melting of the continental
crust.
Chemically, the most distinctive features of the
continental-margin volcanic suites compared to
those erupted in oceanic island arcs are the higher
concentrations of K, Sr, Rb, Ba, Zr, Th and U,
higher KlRb and Fe/Mg ratios and a much wider
range of 87Sr/86Sr, 143Nd/l44Nd and Pb isotopic
compositions. These characteristics must be largely
explained in terms of crustal involvement in the
petrogenesis of the magmas, although the distinc-

tive geochemical characteristics of the subcontinen-
tal mantle wedge may also be important (Section
7.7).
In any destructive plate margin environment
(oceanic or continental) the nature and distribution
of magmatic activity in the overriding plate is
directly linked to the geometry of the subducted
slab (Pilger 1984). This, in turn, is a function of the
convergence rate of the lithospheric plates, the age
of the subducted lithosphere and the presence of
features such as aseismic ridges, oceanic island and
seamount chains, oceanic plateaus and microconti-
nents in the underthrust plate. The latter, by virtue
of their increased crustal thickness, are more
buoyant and tend to resist subduction, frequently
becoming accreted on to the plate margin when
they collide with the landward plate. Numerous
such accreted terrains have now been recognized on
the North American continental margin (Uyeda
1982, Nur &Ben Avraham 1983), making it a more
complex example of an active continental margin
than the South American Andes.
The most complex case of subduction-related
magmatism occurs where two continental plates
approach and collide by subduction of the interven-
ing ocean, e.g. the Alpine-Himalayan system. The
consequent suture zone becomes an area of thick-
ened crust characterized by complex tectonic and
magmatic activity and uplift. After collision, calc-
alkaline andesites and dacites may be erupted,
followed by alkaline volcanism as extensional tec-
tonic regimes develop as a consequence of the rapid
uplift (Harris et al. 1986). Houseman et al. (1981)
have suggested that, during a collision orogeny, the
thickened subcontinental mantle root may become
detached and sink, to be replaced by hot astheno-
spheric mantle which then partially melts as it rises
to produce the post-orogenic magmas. However,
the transition from subduction related to intra-plate
characteristics may not become apparent im-
mediately due to the interaction of the rising
magmas with the hot, thickened continental crust.
The Andean Cordillera of South America ex-
tends for lO 000 km along the western margin of
the continent, from the Caribbean Sea to the Scotia
Sea, making it the longest sub-aerial mountain
chain on Earth. A significant feature of the present-
INTRODUCTION 193
day subduction system is its segmentation into
shallow dipping « 10°) and more steeply dipping
(-30°) zones, with active volcanism occurring only
in association with the steeply dipping segments
(Figs. 7.2 &7.3). This seems at first surprising, as
the rate of convergence of the Nazca plate and the
South American plate is practically uniform (-lO
cm yr-1) along the whole convergence zone (Wortel
1984). The cause of the anomalously shallow
dipping segments has been attributed to the sub-
duction of buoyant aseismic ridges, the Nazca
Ridge and the Juan Fernandez Ridge, within the
Nazca plate. Barazangi & !sacks (1979) have
attributed the absence of active volcanism in north
and central Peru and central Chile (Fig. 7.2), where
the Nazca plate is subducting at shallow angles, to
the displacement of the asthenospheric mantle
wedge and the direct superposition of the two
lithospheric plates (Fig. 7.3). The moderate angle
of subduction (-30°) characteristic of most of the
plate boundary has been attributed to the combined
effects of rapid plate convergence, overriding of the
trench by the South American plate and the relative
youth of the subducting Nazca plate (Cross &
Pilger 1982).
As shown in Figure 7.2, active volcanism within
the Andes is divided into three zones (Thorpe et al.
1982), a northern volcanic zone (NVZ) extending
from SON to 2°S in Colombia and Ecuador, a central
volcanic zone (CVZ) extending from 16°S to 28°S in
southern Peru, northern Chile, Bolivia and Argen-
tina, and a southern volcanic zone (SVZ) in
southern Chile and Argentina. In each of these
zones volcanism has occurred episodically since the
Mesozoic. Table 7.1 summarizes the physical and
geochemical characteristics ofeach ofthese volcanic
zones. The lavas of the NVZ are dominantly
basaltic andesites and andesites, which have miner-
alogical and major element characteristics similar to
island-arc volcanic suites. In general, the lavas of
the SVZ are similar but slightly more basic, with
high-alumina basalt and basaltic andesite being the
most common rock types. The lavas ofthe CVZ are
characteristically intermediate to acid in composi-
tion and show a marked increase in K20 content (at
constant Si02%) with increasing depth to the
Benioff zone, calc-alkaline volcanics grading east-

10'
I
,)
10'
Key
NVZ Northern Volcanic Zone
M Central Volcanic Zone
SVZ Southern Volcanic Zone
• active calc-all<aline
.. I
,----- ,..... I
~ -~ /
volcanoes
A alkaline volcanoes·
l.. , , " " V 5 shoshonitic volcanoes
30'
wards into shoshonites.
o 500
~
km
Andean magmas result from a complex interplay
of partial melting and fractional crystallization
processes within the mantle, and contamination
and fractional crystallization processes within the
crust. Significantly, one of the most obvious
differences between the northern, central and
southern volcanic zones is the occurrence of Pre-
cambrian basement beneath the CVZ, but only
- - oceanic ridges/rises in
0' the Nazca plate
10'
20·
~ _ constructive plale
~ boundaries
,. __, crust> 50 km th,ck and
t __ ,/ rgnimbrite province
O continental areas
unde~aid by crust of
Palaeozoic or younger age
Dzones In which active
volcanism is absent and
uplift and erosion have
revealed granitoid
bathOlith belts
Figure 7.2 Distribution of active volcanoes along the
Andean Cordillera of South America (alter Harmon et
al. (1984). w ith additional data from Thorpe et al.
(1982.1984).
much younger Mesozoic-Cenozoic crust beneath
the NVZ and SVZ. In terms of magma-crust
interaction models (Leeman 1983), it would be
expected that the volumetric proportions oferupted
rock types and their geochemical characteristics
should be strongly correlated with the thickness
and chemical characteristics of the crust through
which the rising magmas have passed. This is
clearly true for the Andes (Section 7.7), supporting

(b) >600km
sea level trench .1 VF
~~~:=~;;~~~~~' I I . I , ,
- ' I I (, I I I , .
<.... ,...fI4i ..... ' .~'/ I
....... _...... '--
Key
~ oceanic lithosphere
lT~~ continental lithosphere
VF volcanic front
Figure 7.3 Schematic illustration of the effect of subduction
of an aseismic ridge on the angle of subduction. (a) A fast
convergence rate and the subduction of young oceanic
lithosphere results in a fairly shallow angle of subduction of
-30°. with the active volcanic front occurring at distances of
150-600 km from the trench. (b) Subduction of the thickened
lithosphere of an aseismic ridge results in the direct super-
position of the continental and oceanic lithospheric plates
over a much greater distance. increasing the distance
between the volcanic front and the trench to >600 km. and in
some instances completely eliminating active volcanism.
(After Cross & Pilger 1982. Fig. 1. p. 547)
petrogenetic models involving interaction of
mantle-derived magmas and their crustal wall rocks
to the extent that some of the more acid magmas
may actually be crustal remelts.
Many Andean volcanic rocks carry a trace
element and isotopic signature of the continental
crust through which they have passed (Section 7.7).
However, as discussed in Chapter 6, destructive
plate margin magmas can also inherit such a
signature from the subduction of continentally
derived terrigenous sediments. Unfortunately, in
most cases it is impossible to separate these two
effects. Opinions are conflicting as to the amounts
of continentally derived detritus currently being
INTRODUCTION 195
subducted beneath the Andes. Shepherd & Mober-
ley (1981) consider that the lack of a substantial
accretionary wedge in the Peru-Chile trench
suggests that all material derived from the continent
has either been subducted or removed laterally,
whereas Uyeda (1982) attributes this to a low rate of
sediment supply to the trench.
One of the characteristic features of the Andean
tectonic setting is the close spatial association of
calc-alkaline volcanic and plutonic rocks, the latter
now generally accepted as the root zones of former
active volcanoes. The intrusive rocks range in
composition from gabbro, though diorite, tonalite
and granodiorite to granite, and show similar
compositional ranges to the volcanic rocks (Section
7.7), strengthening the precept of a genetic rela-
tionship. Collectively, the intermediate to acid
intrusives are known as granitoids.
Studies of subduction-related plutonic rocks
have tended to focus on the western Americas, but
crystalline plutonic rocks of calc-alkaline affinity
also outcrop in many of the island arcs of the
western Pacific, Indonesia, the Aleutians and the
Caribbean. Island arcs do not develop on ancient
continental crust but on a foundation of oceanic
crust, and evolve by thickening of the volcanic pile
by the combined effects of volcanism and pluton-
ism. Eventually, more mature arcs develop con-
tinental crustal-like profiles (Ch. 6) and, during the
Phanerozoic, New Zealand, Japan and Central
America have evolved to this intermediate stage
between immature arcs with thin crust and active
continental margins with Precambrian basements.
The cessation of active volcanism in those
segments of the Andes where the angle of subduc-
tion has decreased to less than 10° is related to
periods of extensive uplift and erosion, revealing
enormous linear batholith belts paralleling the
continental margin. The largest of these is the
Coastal Batholith of Peru, which is over 1600 km
long and 60 km wide and comprises more than 1000
plutons emplaced over a 60 Ma timespan from 100
to 37 Ma. Most of these plutons were emplaced by
permissive cauldon subsidence and stoping within
3-4 km of the surface (Pitcher & Cobbing 1985).
Tonalites and granodiorites (granitoids) predomin-
ate, associated with swarms of basaltic andesite

Table 7.1 Tectonic and geological characteristics of the late Cenozoic volcanic zones of the Andes (after Harmon at al. 1984,
Table 1. p.805).
dip of seismic zone
depth to seismic zone
maximum crustal elevation
crustal thickness
crustal age
composition of volcanics
Si02(wt.%)
K2O(wt.%)
8'SO
87SrJ86Sr
206pbf204Pb
207pbf04Pb
208Pbl204Pb
SVZ
(45-33°S)
<25°
c.90km
2000-4000m
30-35km
Mesozoic-Cenozoic
basalt with minor andesite
and dacite
50-69
0.4-2.8
medium-K series
5.2-6.8
0.7037-0.7044
18.48-18.59
15.58-15.62
38.32-38.51
dykes. The batholith follows a remarkably linear
course over most of its length, suggesting a
fundamental deep control to its emplacement,
postulated to be a deep structural lineament chan-
nelling the magmas upwards. Similar granite bath-
olith belts are exposed along the whole of the
western continental margin of the Americas from
Alaska to the Antarctic, and their emplacement
history spans most of Mesozoic and Tertiary time.
All are intruded within and flanked by subduction-
related volcanic rocks, and many are within regions
of thickened Precambrian crust.
Subduction-related volcanism is one of the main
mechanisms for the growth of the continental crust,
either by lateral accretion of island arcs or by the
vertical addition of intermediate composition intru-
sives and extrusives in active continental margins.
Thorpe eE aI. (1981) have estimated that, for the
Andes, plutonic rocks volumetrically exceed volca-
nic rocks by a factor ofat least 10, attesting to their
fundamental role in crustal growth. Genetically, the
volcanic and plutonic suites must be related,
although many authors have found it difficult to
reconcile the dominantly andesitic composition of
the lavas with the more silicic granitoids of the
batholith.
CVZ
(26-18"S)
c.25-30°
c. 140 km
5000-7000m
50-70km
Precambrian-Palaeozoic
andesite-dacite with
dacite-rhyolite ignimbrites
56-66
1.4-5.4
high-K series
6.8-14.0
0.7054-0.7149
17.38-19.01
15.53-15.68
38.47-39.14
NVZ
(2°S-5°N)
c.20-30°
c. 140 km
4000-6000m
40km
Cretaceous-Cenozoic
basaltic andesite to andesite
53-63
1.4-2.2
medium-K series
6.3-7.7
0.7036-0.7046
18.72-18.99
15.59-15.68
38.46-38.91
7.2 Simplified petrogenetic model
In Section 6.2 it was considered that magma
generation in the subduction-zone environment is a
multistage multisource phenomemon. Much of the
preceding discussion about petrogenetic processes
in island arcs is equally relevant to the active
continental margin environment and therefore will
not be reiterated here. However, in active continen-
tal margins, passage of the magmas through thick
sections of continental crust produces added com-
plexities.
Figure 7.4 shows a schematic cross section of an
active continental margin which might be appropri-
ate for the CVZ of the Andes. Processes in the
subducted oceanic lithosphere should be exactly the
same as those described in Section 6.2 for the case
of ocean-ocean collision. Upon subduction, the
cold oceanic lithosphere is heated by the combined
effects offriction and conduction and, as a consequ-
ence, the oceanic crustal layer undergoes a series of
metamorphic transformations from greenschist
through amphibolite to eclogite facies. Prograde
metamorphic reactions generally involve dehydra-
tion, and the resulting hydrous fluids are released
into the mantle wedge where they lower the solidus

o sea level
50
100
km
150
200
250
and promote partial melting. If the solidus temper-
ature of the subducted crust is exceeded a hydrous
intermediate-acid partial melt may be generated
which similarly metasomatizes the mantle wedge
and causes partial melting. A significant difference
between the model shown in Figure 7.4 and that
shown in Figure 6.3 lies in the thickness of the
continental lithosphere. This is of the order of 140
km beneath the central Andes (Barazangi & Isacks
1979) compared to 70-80 km for typical oceanic
lithosphere (Ch. 5). Additionally, the continental
crust in this profile is 50 km thick, compared to 10
km for average oceanic crust.
In the island-arc tectonic setting (Ch. 6) it is
generally agreed that for volcanism to occur there
must be a wedge of more fertile asthenospheric
mantle above the slab (Gill 1981). In Figure 7.4 a
similar wedge has been shown for the active
continental margin case, although its role must be
considered more equivocal. This is because, unlike
the oceanic lithosphere, which is variably depleted
due to magma generation events at the mid-ocean
ridge, the continental lithosphere may be consider-
ably metasomatized and enriched, particularly if it
has formed part of a stable continental root for a
considerable period of time. Thus slab-derived
fluids could initiate partial melting in the subcon-
tinental lithosphere, adding further complexity to
SIMPLIFIED PETROGENETIC MODEL 197
magma storage reselVoir
Figure 7.4 Schematic
cross section of an active
continental margin.
the isotope and trace element geochemistry of the
magmas. Indeed, Pearce (1983) considers that
enriched subcontinental mantle (lithosphere) plays
a dominant role in the petrogenesis of all basalts
generated in active continental margin tectonic
settings, rather than the convecting asthenosphere.
Any mantle-derived magma passing through a 50
km thick section of continental crust must inevit-
ably interact with it, and therefore assimilation and
fractional crystallization (AFC) processes (Ch. 4)
should be important in the petrogenesis of these
magmas. Of fundamental importance in this re-
spect is the age and geochemical characteristics of
the crust. Ancient Precambrian basement gneisses
have distinctive isotope geochemical signatures
which should readily fingerprint contaminated
magmas (Section 7.7.5), whereas younger
greywacke type sediments may differ little isotopi-
cally from the mantle-derived magmas and there-
fore if these are the crustal contaminant the process
is much harder to identify.
In general, it seems reasonable to assume, as in
Chapter 6, that the primary mantle-derived mag-
mas are basaltic in composition, although genera-
tion of more siliceous magmas from the metasoma-
tized mantle wedge remains a distinct possibility.
Low-pressure crystal fractionation of such magmas,
combined with crustal contamination, can then

account for the spectrum of more evolved rock
types observed at the surface.
7.3. The structure ofactive continental
margms
Four main interdependent variables control the
geometry of subduction zones (Cross & Pilger
1982, Jarrard 1986):
(1) the relative plate convergence rate;
(2) the direction and rate of absolute upper plate
motion;
(3) the age of the subducting plate;
(4) the subduction of aseismic ridges, oceanic
plateau or intra-plate island/seamount chains.
Their effects may be additive, or one variable may
cancel the effect of another.
In general, low-angle subduction results from
combinations of rapid absolute upper plate motion
towards the trench, relatively rapid plate converg-
ence, subduction oflow-density oceanic lithosphere
and the subduction of young oceanic lithosphere.
The consequences are either a landward displace-
ment of the magmatic arc or a cessation of
subduction-related magmatism, and the develop-
ment ofa compressional tectonic regime within and
behind the arc. In contrast, steeper subduction
results from combinations of slow or retrograde
absolute upper plate motion, slow relative rates of
plate convergence and subduction of old dense
oceanic lithosphere. This induces the development
of a magmatic arc closer to the trench and
extensional tectonics within and behind the arc.
Seismology has thus far been the only geophysi-
cal method available for examining the shape,
continuity and physical characteristics of subduc-
tion zones. In Figure 7.5 profiles of the top of the
Benioff zone for three oceanic island arcs are
compared (Marianas, Lesser Antilles and New
Hebrides) with those of three segments of the
Andes in Peru, central Chile and northern Chile.
All of the oceanic examples show considerably
steeper profIles, which may be explained in terms
of the greater age of the subducting plate and the
100
200
E 300
~
.s::.
g. 400
o
500
600
700
trench
Distance from trench (km)
(a) Island ares
Marianas
Distance from trench (km)
O~~r-~~~--~~-r__~__~~
E
~
100
200
-5 300
c.
al
o
400
500
600
(bl Andean margins
Figure 7.5 Profiles of the top of the Benioff zone for three
intra-oceanic island arcs (a) compared with three segments of
the Andes (b). The vertical dashed line indicates the location
of the volcanic front in each case (after Jarrard 1986. Fig. 2. p.
223).
slow rate of plate convergence. Comparing the
three Andean profIles, those from Peru and central
Chile are characterized by a very shallow dip
(<100) and lack active volcanism. These are among
the most shallow dipping of all modem subduction
zones Garrard 1986) in which the slab appears to
flatten and travel horizontally along the lower
surface of the South American lithosphere. The
shallow angle of subduction has been attributed to
the subduction of buoyant aseismic ridges in the

STRUCTURE OF ACTIVE CONTINENTAL MARGINS 199
underthrust Nazca plate (Cross & Pilger 1982). In
contrast, the moderate angle of subduction (-30°)
characteristic of much of the plate margin may be
attributed to the combined effects of rapid converg-
ence, overriding of the trench by the South
American plate and the relative youth of the
subducting Nazca plate.
During the evolution of the Andean subduction
system over the past 250 Ma there have doubtless
been numerous changes in the angle of subduction
and the consequent location of the volcanic front.
For example, between 50 and 25 Ma the age of the
subducting oceanic plate progressively decreased
(Pilger 1981), resulting in a gradual decrease in the
subduction angle and consequent eastward shift in
the loci of magmatic activity. Additionally, a major
reorganization of plate motions at 25 Ma resulted in
an increased convergence rate normal to the Chi-
lean Andes, with a consequent shallowing of the
subduction zone.
As suggested in the introduction, Andean mag-
matism results from a complex interplay of crystaV
melt equilibria in the mantle, and contamination
and fractional crystallization processes in the crust.
To understand the nature of the crustal contamina-
tion process it is important to have as much detailed
information as possible about the crustal structure.
Seismic refraction surveys provide information
about the thickness and velocity structure and
enable identification of upper and lower crustal
layers (Fig. 7.6). In this diagram crustal profiles for
Andean segments in Colombia and northern Chile
are compared with profiles for the Cascades, Alaska
and New Zealand and for Tonga, an immature
island arc. It is important to realize that the seismic
data can only give information about the broad
lithology of the crustal profile, not the age. Thus
lower crustal rocks with a Vp of 6.5-7.0 beneath
Tonga are undoubtedly much younger than those
with a similar range of Vp beneath northern Chile.
Upper crustal rocks with Vp 5.0-6.2 comprise
sedimentary rocks, volcanics and young
intermediate-acid plutonic rocks. In contrast, the
lower crust with Vp of 6.5-7.4 may be considered
to be made up of high-grade crystalline metamor-
phic rocks in the granulite and amphibolite facies.
Continental Continental Island
arc
o
10
20
30
<m
40
50
60
70
Continental margins
Colombia
77-
7.9
N. Chile Cascades
7.8
peninsula fragment
Alaska New Zealand Tonga
....
....
....
.,. ..
t •••
,. ...
8 1 8.0
Vp (km 5- ')
7.7
t8m 5.0-5.7
upper crust ~
~ 58-6.2
{
Y77/I
~6. 5-70
lower crust ~
~72-74
mantle D >77
Figure 7.6 P-wave
velocity (Vp) crustal profiles
for a variety of active
continental margin tectonic
settings. compared with
that for an immature intra-
oceanic island arc (after Gill
1981. Fig. 33. p. 48)

trench
Western
Cordillera
Eastern
Cordillera
Brazilian
Platform
o
~. I _ - .
..... -.._' -" ' ' - / .,"
..... - - , I ...,
. - - -- ~. 'r,'. '- .. :
( I , - ..
'"" """ "' ,'~O,,O
., . ~._' -: .7-: <:>',' .!
- - :: ' ',,- ..~ ,
/ ...... 1 : ~ ,.......
.... • _~ I
100
!11
o 400 600
Distance (km)
Figure 7.7 depicts the crustal structure of the
central Andean plate margin. This shows that the
bulk of the upper crust in the. active volcanic belt
comprises a young granitoid batholith overlain by
intermediate composition volcanic rocks. The high-
grade ancient metamorphic rocks of the Brazilian
shield pass beneath the batholith and outcrop on
the Pacific coast. Clearly, in this region of thick-
ened continental crust, magmas must pass through
in excess of 50 km of high-grade Precambrian
gneiss before reaching the high-level magma stor-
age reservoirs. Therefore, it is not in the least
surprising that the CVZ magmas carry a strong
imprint of continental crustal contamination in
comparison to magmas generated in the NVZ and
SVZ, where ancient gneissic basement complexes
are absent.
Figure 7,8 is a schematic cross section of the
Peru-Chile subduction zone, showing the dis-
tribution of zones of moderate and high seismicity
and the variation of the seismic attenuation factor,
Q. The upper zone of the subducted oceanic
lithosphere is characterized by high seismicity and
high Q, while the overlying mantle wedge is
Key
~
v volcanics
:~":~:~ :'~ sediments
intruSives
I~'<. shlekJ gneiss
I I
800 1000
Figure 7.7 Schematic
cross section through the
central Andean active
continental margin to show
the crustal structure (x 5
vertical exaggeration).
Arrows indicate the
direction of magma and
volatile streaming from the
downgoing plate (after
Brown & Mussett 1981 .
Fig. 96. p. 168).
trench km
200 0 200 400 600 800 1000
100
200
' ..
300 ....: . •: :.... :~...::':
km . .. . ' ..,.
..
. . . " . ' -'
400 . '.: : ".:..: ........ ',,:- .:'..... .
500
600
..' .. .
" , .
. : ' ,' " : . - . ... -,' .... :',:::.:: : ':' ...... '
: : ~'...:..:.. ,~.~-!......:.. . ---. ~:.......:.. .~.:...:.:.. ~ .~.~, ,.'
o
~ moderate
I888l high
Seismiciry
Figure 7.S Schematic cross section through the Peru-Chile
subduction zone. showing the distribution of zones of
moderate and high seismicity and the variation of the seismic
attenuation factor 0 (after Condie 1982. Fig . 6.
12. p. 112).

THERMAL STRUCTURE AND PARTIAL MELTING PROCESSES 201
moderately seismically active and also has high Q
(1000-3000). Sacks (1983) suggests that this indi-
cates the existence of a thickened zone of subcon-
tinentallithosphere (-350 km) above the slab. Low
values of Q «200) indicate high seismic attenua-
tion and the probable presence of a partial melt
phase. Below a depth of about 400 km the
subducting slab appears to flatten and to develop
low Qcharacteristics suggestive of partial melting.
7.4 Thermal structure and partial
melting processes
The thermal structures of subduction zones are not
well defined (Section 6.5) and a variety of models
have been proposed (Anderson et al. 1978, 1980;
Furlong et al. 1982). The subducted oceanic crust
may be relatively warm as a consequence of
frictional heating, or it may be significantly cooled
by endothermic dehydration reactions. Similarly,
the mantle wedge may be relatively cool, chilled by
the subducting slab, or it may be heated by induced
convection (Toksoz & Hsui 1978, Wyllie 1984).
Figure 7.9 shows the schematic thermal structure
of an active continental margin with a SO km thick
o
50
km
100
150
(al Cool mantle wedge, chilled by the
subducting slab
-------
I
{
I

f:§:Jr:;S-
,,'"
crust for two extreme thermal models which may be
described as cool (A) and warm (B) respectively.
The consequence of induced convection in the
mantle wedge (model B) is to raise the 750°C
isotherm into the base of the continental crust,
thereby increasing the likelihood of crustal melting.
Additionally, the 1000° and 12S0°C isotherms are
raised higher in the mantle and thus this latter
thermal regime will be more conducive to extensive
mantle partial melting. Figure 7.10 shows the same
thermal models with the approximate location of
the solidi in the presence of H20 for the major
sources which may contribute to partial melts in the
subduction-zone environment; the subducted
oceanic crust, the continental crust and the mantle
wedge. The line D-D' marks the onset of signifi-
cant dehydration within the slab, broadly coinci-
dent with the greenschist/amphibolite facies bound-
ary. Hydrous fluids are shown streaming into the
mantle wedge and the base of the continental crust
where, under subsolidus conditions, they may
promote the growth of extensive metasomatic
amphibole (hatched regions). The dotted line in
each diagram shows the maximum depth of stabil-
ity of amphibole. Partial melting occurs where any
of these three major sources are raised above their
(bl Warm mantle wedge, heated by
induced connection
Figure 7.9 Schematic thermal structure of an active continental margin (after Wyllie 1984. Fig. 8, p.449).

E
~
'5.
Q)
o
o
50
100
150
o
50
100
150
la) Cool mantia, cool subducted crust
CONTINENTAL CRUST
solidus - crust
Ib) Warm mantle, warm subducted crust
OCEANIC CfI
r---_. Us.,.
CONTINENTAL CRUST
solidus - crust
solidus - mantle
AMPHIBOLITE
Key
partial melt
subsolidus metasomatic
growth of amphibole
D-D' dehydration front for
subducted oceanic crust
solidi in the presence
of H20
Figure 7.10 Location of
the sites of partial melting
in active continental
margins for two different
thermal models (after
Wyllie 1984, Fig. 9, p.450).

respective solidus temperature, particularly where
H20 streams into regions to the high-temperature
side of the solidus.
For both thermal models, aqueous fluids pene-
trating the base of the continental crust may
promote partial melting. However, only in thermal
model B does extensive mantle partial melting
occur, triggered by partial melts ascending from the
subducted oceanic crust. These mantle partial
melts then rise into the base of the continental
crust, where mixing with anatectic crustal melts
may occur. Model B may be considered to be
generally applicable to the Andean tectonic setting,
and thus we can see that the erupted volcanics may
contain a contribution from each of the three
potential magma sources. If the subducted oceanic
crust contains a significant proportion of continen-
tally derived terrigenous sediment, it may be
difficult to resolve geochemically components with
a continental crustal signature inherited from the
subduction zone from those introduced by high-
level crustal contamination. Oxygen isotope studies
may be particularly useful in this respect (Section
7.7.6).
Many workers have suggested the importance of
crustal melting in the generation of the vast sheets
15.-------",---------------,
~
~
i' H20-UNDERSATURATED LIQUIDS
Temperature (OC)
Figure 7.11 Compositions of liquids generated by partial
melting of continental gneisses in the presence of H20 (2%).
The shaded field depicts the conditions under which H20-
saturated liquids can occur (after Wyllie 1984, Fig. 6, p.445).
MAGMA STORAGE IN THE CRUST 203
ofignimbrite characteristic ofthe CVZ ofthe Andes
(Gill 1981). Figure 7.11 shows the range of
compositions of partial melts which might be
derived from continental gneisses in the presence of
H20 (Wyllie 1984). There is clearly a very narrow
temperature interval for the existence of H20-
saturated rhyolitic liquids close to the solidus,
except for pressures less than 2 kbar, and at all
crustal depths very high temperatures are required
to derive andesitic liquid compositions. Increasing
pressure produces liquids with lower Si02 contents
and at the base of a 40-SO km thick crust the
near-solidus partial melt may be syenitic (Huang &
Wyllie 1981). Thus it is possible that a range of the
more acidic Andean magma compositions could be
generated by direct partial melting of the continen-
tal crust.
7.5 Magma storage in the crust
Evidence for the existence of shallow magma
reservoirs in the crust beneath active volcanoes is
provided by the following:
(1) geophysical data;
(2) petrological evidence for the role of low-
pressure crystal fractionation in the geo-
chemical evolution of the magmas;
(3) the existence of plutons underlying eroded
volcanic complexes.
Geophysical techniques for detecting magma
bodies are based on the dramatic decrease in
density and seismic velocity, and increase in
seismic attenuation and electrical conductivity,
which occur at the onset of partial melting in rocks.
Seismicity beneath volcanoes is caused by magma-
induced tectonic stresses, and if extensive zones of
partial melt are present (Le. magma chambers)
earthquake stresses cannot accumulate. Thus zones
of seismic quiescence may indicate the locations of
crustal magma reservoirs. Iyer (1984) has reviewed
the geophysical evidence for the locations, shapes,
sizes and internal structure of magma bodies
beneath selected regions of Quaternary volcanism
including Alaska and Kamchatka (continental

peninsulas) and New Zealand (continental frag-
ment), all of which may broadly be considered as
examples of active continental margins. Unfortu-
nately, these data are limited and similar high-
quality seismic data for the Andean margin are
lacking altogether. For Kamchatka there is evi-
dence for magma storage bodies in the depth range
30-90 km, with dimensions between 8 and 40 km
across and up to 30 km thick. However, in some
instances much shallower reservoirs occur, within
10 km of the surface, fed by conduits extending into
deep-seated mantle magma reservoirs. Marked
low-velocity zones have been recorded at depths of
10 and 35 km in the crust of the central Andes, and
have been interpreted by Ocala & Myer (1972) as
potential zones of magma storage.
In active continental margins, volcanic and
plutonic rocks, ranging in composition from basalt
(gabbro) to rhyolite (granite), frequently display
good linear correlations on Harker variation dia-
grams (Section 7.7), suggestive of the derivation of
the more acid magmas by fractional crystallization
of olivine, plagioclase, pyroxene, magnetite and
amphibole mineral assemblages from basaltic
parent magmas. In suites of volcanic rocks for
which Sr-Nd-Pb isotopic data suggest little
crustal contamination, these data may be inter-
preted as reflecting liquid lines of descent. Howev-
er, most Andean magmas, particularly those
erupted in the CVZ, have geochemical characteris-
tics reflecting the combined processes of assimila-
tion and fractional crystallization (Ch. 4) which, in
general, will tend to blur coherent linear trends on
Harker diagrams. Nevertheless, there still appears
to be abundant geochemical evidence for low-
pressure crystal fractionation trends, providing
strong supporting evidence for the existence of
high-level crustal magma reservoirs.
One of the most useful lines of evidence in
elucidating the structure of high-level magma
chambers beneath active volcanoes is to examine
the plutonic root zones of deeply dissected volcanic
belts, the granitoid batholiths. Previous to the
1970s, the plutonic and volcanic phases of Andean
magmatism tended to be treated as separate uncon-
nected phenomena, based on the incorrect assump-
tion that the batholiths were dominantly granitic as
opposed to the intermediate composition of the
volcanic belt. However, it is now well established
that the study of individual plutons comprising the
batholith can provide invaluable information about
high-level « 10 km) magma storage reservoirs.
Mesozoic and Cenozoic batholiths are exposed in
the mobile belts of the western Americas, attesting
to the continuity of subduction-related magmatism
along the whole of the continental margin from the
late Triassic. The Coastal Batholith of Peru is some
1600 km long by 60 km wide and up to 15 km
thick, elongated parallel to the present trench. It is
composed of over 1000 plutons intruded over a 70
Ma period from 100 to 30 Ma. The plutonic rocks
are spatially coincident with two groups of volcanic
rocks, the 100 Ma Casma group and the early
Tertiary Calipuy group which overlies an erosion
surface cut through the batholith. The batholith is
divided into five segments (Fig. 7.12), exhibiting
recognizably distinct groups of plutonic rocks
which may be related to discontinuities in the
underlying subduction system at their time of
formation, similar to the segmentation of the
presently active volcanic zone.
The plutonic rocks ofthe batholith comprise 16%
by volume gabbro and diorite, 58% tonalite and
granodiorite, 25.5% adamellite and 0.5% granite
(Hughes 1982). This clearly indicates that the term
'granite batholith' is a misnomer as intermediate
composition rocks predominate. For much of its
length the coastal batholith occupies the axis of an
early Cretaceous marginal basin (Pitcher et al.
1985), although to the south it penetrates old
crystalline basement. The batholithic magmas
appear to have been channelled along the same
deep-seated suture along which the marginal basin
opened. Magmatism within the batholith was
distinctly episodic (Fig. 7.13), with quiescent
periods often longer than 15 Ma between intrusive
phases (Beckinsale et al. 1985).
At the present rather shallow «5 km) level of
erosion the batholith comprises arrays of intersect-
ing plutons, forming complexes with a surprisingly
regular spacing of 120 km. Pitcher et al. (1985)
relate this 120 km spacing to the location ofseparate
melt cells at depth. At such high crustal levels
magmas are hydrostatically emplaced by a com-

o · ("
. .J •
~~. j
C/ c:::...-.
~ D~~..
J

MAGMA STORAGE IN THE CRUST 205
~
o 300 ~.m
,
'--------',
bination of roof lifting and cauldron subsidence
(Pitcher 1979, Pitcher et ai. 1985) and the shapes of
plutons are controlled by magma-induced fracture
patterns (Bussell 1976). In three dimensions, the
roofs of the plutons are flat with rapid turn-downs
into steep sides, forming a box-like shape. Such
plutons may be simplistically represented as the
magmatic filling of a cavity above a down-dropped
block of pre-existing country rock. The majority of
plutons have a circular outcrop pattern, with some
degree of elongation along the structural grain of
(-.... .)
. />
,-..
t.__.
:~
Abancay ~.~
- -a ••
"
~

(
<
'7
)
the country rocks.
Figure 7.12 The
segmentation of the
Cretaceous Coastal
Batholith of Peru (shown
by various ornaments or
blank): Also shown in black
is ~ belt of Cenozoic stocks
and batholiths paralleling.
but to the landward side of.
the Coastal Batholith (after
Pitcher & Cobbing 1985.
Fig. 3.2. p. 22).
Figure 7.14 shows a schematic cross section of
the batholith showing the nested belljar like plutons
intruded into a basement of pre-Cretaceous rocks
overlain by volcanics of the 100 Ma Casma group,
which may represent the earliest phase of volcanic
activity associated with the oldest plutons of the
batholith. The cross section shows the youngest
phases of plutonism venting to the surface to
produce the Tertiary volcanic cover of the Calipuy
group (Cobbing et ai. 1981).

o
50
Age (M a)
'"
:>
o
'"
u
100 !9
150
200
~
u
TOQuepala
segment
Arequ ipa
segment
lima
segmenl
Figure 7.13 Major intrusive phases of the Toquepala,
Arequipa and Lima segments of the Coastal Batholith of Peru
(after Beckinsale et al. 1985, Fig. 16.10, p. 198).
Key
~Ca~as }
rnSavan
mSan Jer6nimo
Dpuscao
monzogranites
~ Huampi Piruroc granodiOrite
EmSan.a Rosa granodiori.e
~ San.a Rosa lonali.e
(=J~{1J Paccho quartz d;orit&-tonallte
7.6 Petrographic characteristics of the
volcanic and plutonic rocks
In Section 6.10 the petrographic characteristics of
the four major magma series erupted in oceanic
island arcs (tholeiitic, calc-alkaline, high-K calc-
alkaline and shoshonitic) were described. Chemi-
cally similar magmas erupted in active continental
margin tectonic settings are virtually identical and
thus the reader is referred to Section 6.10 for the
relevant information, and to Ewart (1982) for a
more detailed synthesis. In this section emphasis is
placed on the petrography of plutonic rocks of the
calc-alkaline series, as these form the bulk of the
exposed granitoid batholith belts, Much of this
information is also relevant to the study of
subduction-related plutonic rock associations ex-
posed in the more mature oceanic island-arc
systems.
Figure 7.15 shows the distribution of the major
rock-forming minerals in rocks, ranging in com-
position from gabbro to granite from the Coastal
Batholith of Peru (Mason 1985). The mineralogy
and textures of these rocks reflects a history of
magmatic crystallization in high-level subvolcanic
magma chambers. However, as with all slowly
cooled plutonic rocks there is abundant evidence
• Palap gabbro-dioflle
. volcaniCS
~CahPYV group} .
~ ~ ~ }Casma group
ff~;!' Pfe-Cretaceous roc1cs
o
I
10,m
I
Figure 7.14 Cross section of the Coastal Batholith of Peru, showing the nested belljar-shaped plutons. PS is the trace of the
present topography (after Bussell & Pitcher 1985, Fig. 15.4, p. 169).

Gabbro Diorite Tonalite
Grano-
Granite
diorite
Iolivine
Icpx
Ipigeonite
lopx I I I I
Iamphibole II II II IIII I
Ibiotite II I
Imagnetite II IIII IIII
Iplagioclase
Ialkali feldspar I
Iquartz I
Figure 7.15 Distribution of the major rock-forming minerals
in calc-alkaline plutonic rock suites (after Mason 1985),
for the growth of subsolidus minerals such as
biotite, amphibole and chlorite due to the inter-
action of the solid rocks with high-temperature
hydrothermal fluids. Figure 7.16 shows some of the
characteristic textural features of a calc-alkaline
granodioritic plutonic rock.
The major rock-forming minerals are plagio-
clase, alkali feldspar, quartz, pyroxene, amphibole,
biotite and magnetite. Sphene and apatite are
common accessory minerals, even in the more basic
rocks, while allanite occurs quite frequently in the
highly differentiated granites.
Pyroxene. The dominant pyroxene phase is an
augite or calcic augite, joined by hypersthene in
the intermediate composition range. Inverted
pigeonite occurs in some of the gabbros, and
Mason (1985) suggest that it may be a high-
pressure phenocryst phase in the more basic
magmas. The occurrence of calcic augite and
hypersthene is considered to reflect relatively
high water fugacities during crystallization.
PETROGRAPHIC CHARACTERISTICS OF ROCKS 207
Figure 7.16 Characteristic textural features of a calc-alkaline
granodiorite from Chile, (a) Multiply twinned plagioclase with
interstitial quartz and k-feldspar (x40, crossed polars), (b)
Intergowth of amphibole, biotite and magnetite (x40, ordin-
ary light).
Amphibole. Hornblende is one of the major mafic
minerals crystallizing from magmas ranging from
basic to acid in composition. This is in marked
contrast to its occurrence in calc-alkaline volcanic
suites, in which it occurs infrequently and often
in a highly resorbed state. The abundance of
hornblende in the plutonic rocks reflects the
increased stability of amphibole at depth in the
crust. Crystals are generally euhedral or subhed-
ral, indicating early crystallization, and change in
colour from brown through green-brown to
green with increasing differentiation of the mag-
ma. The colour changes appear to correlate with
progressively decreasing Ti02 contents. In some

rocks an original green- brown hornblende may
be patchily replaced by green hornblende and
associated sphene. This is most probably a
solid-state reaction product in the presence of a
hydrothermal fluid phase. Early formed amphi-
boles in the basic rocks are tschermakitic horn-
blendes, whereas in the acid rocks later formed
amphiboles tend to be actinolitic hornblendes.
Biotite. Biotite is a common mafic mineral in
many granitoid rock types, appearing late in the
crystallization sequence of the more basic rocks
but early in the more acid intrusives, where it
may form well developed crystals. MgI(Mg +
Fe2+) ratios vary from 0.38 to 0.61 proportional
to those in the host rock. Biotite may be quite
commonly altered to chlorite as a consequence of
interaction with late-stage hydrothermal fluids.
Plagioclase. Plagioclase is the major rock-forming
mineral in nearly all the plutonic rocks, ranging
in composition from An93 to AnlO' The crystals
often show complex oscillatory zoning similar to
that observed in plagioclase phenocrysts in
andesitic lavas. This is a characteristic feature of
the intermediate to acid rocks. Fine-scale
myrmekite (plagioclase-quartz intergrowth) is
common in all rock compositions, but particular-
ly so in the more basic rocks.
Alkali feldspar. The amount of alkali feldspar
present in the plutonic rocks varies in a regular
manner with the bulk rock composition. In more
basic rocks it tends to occur interstitially, where-
as in the more acid rocks it forms larger 'pools'.
Some of the granitoids contain K-feldspar mega-
crysts which are generally considered to have
been producted by late-stage K-rich metasomat-
ism (subsolidus). Orthoclase is by far the most
common type of K-feldspar in the granitoids,
while microcline occurs only in some of the most
differentiated rocks. The degree of ordering in
the K-feldspar seems to be mainly controlled by
the concentration of volatile components in the
melt, with microcline crystallization being
favoured by the most volatile-rich conditions.
Exsolution textures are ubiquitous, although the
alkali feldspar observed in basic rocks normally
lacks exsolution lamellae and is probably a
cryptoperthite. Vein perthites are dominant in
the intermediate and acid rocks, while patch
perthites are most common in the most evolved
rocks. Parsons (1978) has suggested that magma-
tic water might be the prime catalyst in causing
perthite coarsening. Granophyric intergrowths
are characteristic of the most highly differenti-
ated rocks which formed from the most volatile-
rich magmas. These are considered to have
formed from the rapid crystallization of quartz
and alkali feldspar as a consequence of a sudden
reduction in vapour pressure due to loss of
volatiles from the system (Mason 1985).
Magnetite. Magnetite is the major opaque oxide
phase throughout the spectrum of basic to acid
magmas, with ilmenite occurring only rarely.
Both phases tend to exhibit high degrees of
subsolidus re-equilibration.
7.7 Chemical composition ofthe
magmas
7.7.1 Charactistic magma series
The four major magma series recognized in oceanic
island arcs (low-K, calc-alkaline, high-K calc-
alkaline and shoshonitic; see Section 6.7) also occur
in active continental margin tectonic settings. Their
classification is based upon the same KzO versus
SiOz diagram, and the reader is referred to Section
6.7 for further details. However, in comparison
with island-arc volcanic suites (Fig. 7.17), low-K
series magmas are poorly represented, while high-
K and shoshonitic magmas are more common,
particularly at the acid end of the spectrum. These
high-K characteristics may reflect increasing de-
grees of crustal contamination in the active margin
magmas. Additionally, suites of alkaline volcanic
rocks may occur to the landward side of the
volcanic front, ranging from mildly alkaline basalts
to leucite basanites and their derivatives. These
magmas are not necessarily subduction-related, and
may be generated as a consequence of extensional
tectonics in a back-arc region.

CHEMICAL COMPOSITION OF THE MAGMAS 209
Frequency
la) Ib)
basalts.
<52%
Si0 2
basaltic
andesites.
52- 56%
SiO,
andesite.
5~3%
SiO,
dacite.
63-69%
Si0 2
rhyolite.
>69%
Si0 2
Figure 7.17 Comparison of the relative frequency of occurrence of rocks of the low-K. calc-alkaline (CAl. high-K calc-alkaline
and shoshonitic (5) series in ('!) the Andes and (b) the oceanic island arcs of the south-west Pacific (data from Ewart 1982.)
Figure 7.18 compares the frequency distribution
of basalts, basaltic andesites, andesites, dacites and
rhyolites, irrespective of magma series, in the
Andes with that in the island arcs of the south-west
Pacific (Ewart 1982). This clearly reveals the
greater abundance of intermediate and acid mag-
mas erupted in the active continental margin
tectonic setting which, as stated previously, may be
a consequence of crustal contamination.
7.7.2 Major elements
Si02, Ti02, A120 3, Fe203, FeO, MnO, MgO,
CaO, Na20, K20, P20 S and H20 can all be

60
{al
40
20
Or-~B--~----'-----'---~-'--R--~
%
60
(bl
40
20
Or-----,-----,-----,------.----4
8 R
Figure 7.18 Frequency distribution of basalts (B). basaltic
andesites (BA), andesites (A). dacites (D) and rhyolites (R) in
the Andes (a) compared with that in the island arcs of the
south-west Pacific (b) (Data from Ewart 1982.)
considered as major elements in the description of
the geochemistry of active continental margin
magmas. In terms of these, the most obvious
distinction between the major magma series is one
of increasing total alkali content in the sequence
tholeiitic - calc-alkaline - high-K calc-alkaline -
shoshonitic, K20 showing proportionately the grea-
ter increase. This has already been used in Section
6.7 as the basis for the classification of island-arc
volcanic suites.
Figure 7.19 is a plot of wt. % K20 versus wt. %
Si02 for recent volcanic rocks from the northern
(NVZ), central (CVZ) and southern (SVZ) zones of
the Andes. Volcanics from the NVZ and SVZ have
medium-K or calc-alkaline characteristics and are
restricted to Si02 contents <63% (i.e. dacites and
rhyolites are lacking). In contrast, volcanics from
the CVZ have generally high-K characteristics,
spanning the complete compositional range from
basalt to rhyolite. Figure 7.20 is a comparable plot
for plutonic rocks from the Arequipa and Lima
segments of the Coastal Batholith of Peru, showing
that there is total overlap between the compositions
of Andean volcanic and plutonic rocks. In island-
arc volcanic suites K20 behaves essentially incom-
patibly, and thus genetically related suites of rocks
define positive linear trends in plots of K20 versus
Si02 (Section 6.7). While this is also broadly true
for the active continental margin volcanic and
plutonic suites, there is a considerably greater
degree of scatter which may be attributable to the
effects of crustal contamination.
Suites of rocks related by fractional crystalliza-
tion processes and unmodified by extensive crustal
contamination should also define coherent linear
trends on all types of Harker diagram. For exam-
ple, Figure 7.21 shows the variation ofwt.% MgO,
CaO and Alz03 versus % Si02 for plutonic rocks
from the Lima and Arequipa segments of the
Coastal Batholith of Peru. The data define remark-
ably good linear trends, consistent with the frac-
tionation of ferromagnesian minerals and plagio-
clase from parental basalts, bearing in mind the
difficulty of obtaining true liquid compositions by
analysis of plutonic rocks because of the effects of
crystal accumulation. However, caution must be
exercised in interpreting such trends as true liquid
lines of descent until the isotopic homogeneity of all
members of the plutonic suites is verified (Section
7.7.5). Shown for comparison in Figure 7.22 is a
plot of wt. % MgO and wt.% K20 versus wt.% Si02
for volcanic rocks from the Tertiary Calipuy group
of Peru, which overlies an erosion surface cut
through the Coastal Batholith. Coherent trends are
still visible, although they are rather more noisy
than the plutonic data. Both volcanic and plutonic
suites show typical calc-alkaline differentiation
trends with total iron content decreasing progres-
sively as the Si02content increases due to the early
crystallization of magnetite.
Table 7.2 shows average major element analyses
of Andean volcanic rocks compared to those from
the island arcs of the south-west Pacific. These data
clearly show that magmas erupted in the Andean
region are enriched in K20, Na20, Ti02and P20S
and depleted in CaO, compared to their island-arc

4.0
3.6
3.2
2.8
2.4
•
•
0.8
0
.,;
~
0
~
4.0
IN. Chile, NW. Argentina,
3.6
SW. Bolivia)
•
• •
3.2
2.8
2.4
•
•• ••
• ••
0.8
51
••
•
•
•
•
•
•
high-K
•
med-K
high-K
med-K
low-K
•
IS. Peru)
Wt % Si02
•
•
I.
•••
•
•
high-K
med-K
low-K
high-K
• •
med-K
Figure 7.19 Plots of wt,% K20 versus wt.% Si02 for young volcanic rocks from the northern. central and southern volcanic
zones of the Andes. The boundaries between the low-. medium- and high-K fields are those of Peccerillo &Taylor (1976) (after
Harmon et ai, 1984. Fig. 2. p. 810),
counterparts (Ewart 1982). In Table 7.3 average
compositions of basaltic andesites (52-56% Si02)
from the NVZ, CVZ and SVZ of the Andes are
compared. The CVZ basaltic andesites appear
slightly richer in Ti02 and K20 than those from the
NVZ and SVZ, but otherwise the analyses are
broadly similar. In Table 7.4 a typical basalt from
the SVZ is compared with an alkali basalt erupted
in an extensional tectonic setting to the east of the
volcanic front in the CVZ. The alkali basalt is much
poorer in Si02 and therefore the two analyses are
not directly comparable. However, it is evident that
the alkali basalt has much higher concentrations of
TiOz and P20 S and the whole range of incompati-
ble trace elements and lower A120 3 • Table 7.5
shows whole-rock analyses of plutonic rocks from
the Lima segment of the Coastal Batholith of Peru,
for comparison with the volcanic data.
7.7.3 Trace elements
In Section 6.7 it was demonstrated that island-arc
basalts are characterized by selective enrichment of
elements of low ionic potential (Sr, K, Rb, Ba ±

5.0
Key
4.5
• Arequipa segment high-K
o LIma segment
4.0 o
•
3.5
0
3.0
~
<f. 2.5
o med-K
0 o
o
.0 •
:
ce
10 I
I.
0
0.51 I
49 51 53 55 57 59 61 63 65 67 69 71 73 75
Figure 7.20 Plot of %
K20 versus O/OSi02 for
plutonic rocks from the
Arequipa and Lima
segments of the Coastal
Batholith of Peru (data
from Pitcher et al. 1985).
% Si02
Th) and low abundances of elements of high ionic
potential eTa, Nb, Ce, P, Zr, Hf, Sm, Ti, Y, Vb, Sc
and Cr) compared to N-type MORB. The enrich-
ment in low ionic potential elements has been
attributed to metasomatism of the mantle source of
arc basalts by fluids released from the subducted
slab. In contrast, the relative depletion in high ionic
potential elements has been variably attributed to
higher degrees of partial melting and to the stability
of residual mantle phases (Pearce 1982).
Figure 7.23 shows chondrite-normalized trace
element abundance patterns (spiderdiagrams) for
basaltic andesites from the northern, central and
southern volcanic zones of the Andes. More primi-
tive basaltic compositions would normally be used
for such a diagram, but unfortunately data are
unavailable. Compared to the equivalent diagram
for oceanic island-arc basalts (Fig. 6.34), they
clearly show the same distinctive spiked pattern
with peaks at K, Sr and Th and a marked trough at
Nb. It appears that such patterns must be a
characteristic of all subduction-related magmas,
attesting to the involvement of subduction-zone
fluids enriched in Sr, K, Rb, Ba and Th in their
petrogenesis.
Figure 7.24 shows a MORB-normalized trace
element variation diagram (Pearce 1983) for the
least enriched of the two CVZ basaltic andesites
shown in Figure 7.23. Comparing this with the
patterns for intra-plate and island-arc basalts in
Figure 6.37, we can see that the immobile elements
Ta, Nb, Zr, Hf, Ti, Y and Yb define a pattern
(dashed line) more akin to that ofintra-plate basalts
than to MORB. Following Pearce (1983) it is
suggested therefore that the mantle source of this
magma was enriched subcontinental lithosphere (as
opposed to depleted asthenosphere in the case of
island-arc basalts) to which mobile elements (Sr, K,
Rb, Ba, and to a lesser extent Ce and Sm) had been
added by a subduction-zone fluid. Shown for
comparison in Figure 7.24 is the trace element
pattern for an alkali basalt erupted to the east of the
CVZ in an extensional tectonic regime. This shows
a typical intra-plate signature (Fig. 6.37) and may
be derived by partial melting of the subcontinental
lithosphere, possibly with some contamination by
the continental crust. Thus Figure 7.24 clearly
attests to the involvement of subcontinental litho-
sphere as a major source component in the
petrogenesis of Andean volcanic rocks.
A difficult question to resolve is how patterns
such as those in Figure 7.24 reflect involvement of

20
18
16
14
8
•
6
4
6
(a) Lima segment
•
•
• •
• • • ... AI,0 3
·' ..... ..
...... ..
.. .:.~
. ..,
,...
••,.1.
•
• •
.~..~ .
• ... •• ~.. CaO
. . ..,
. ·'1
•• , .-4.
..., ' ••4l
...••
CHEMICAL COMPOSITION OF THE MAGMAS
20
18
10
8
..
6
4
0
6
4 •
(b) Arequipa segment
• • • •
,, • ••
• • .:••
• •
•
•
•
• • ,... ...
• • • •
• • ••
• •• •
• •••
.. : .,...
•
, •
•
• •
•
•
•
I •• •
213
AI20 3
CaO
MgO
°5~5------~~------L-------~------~75
% Si02
Figure 7.21 Variation of A120 3• CaO and MgO versus Si02 for plutonic rocks from the Lima and Arequipa segments of the
Coastal Batholith of Peru (data from Pitcher et al. 1985).
continental crustal materials. This is particularly
important in the petrogenesis of the CVZ Andean
magmas, as we shall see in Section 7.7.5. The trace
element signature of crustal contamination is parti-
cularly difficult to predict, given the great range of
crustal rocks which could be involved and the
likelihood that the contaminant will be a partial
melt of one of these rocks rather than the bulk rock
itself. Figure 7.25 illustrates the types of trace
element patterns that might result from selective
contamination of a basalt (with MORB-normalized
abundances of0.5) with 50% partial melts ofdiorite
and greywacke crustal rocks respectively, in the
proportion 4: 1, basalt: contaminant (Pearce
1983). This is obviously an extreme case, as the
addition of such a large volume of acidic partial
melt would change the composition of the basalt to
that of a basaltic andesite or andesite. Of particular
significance is the fact that crustal contamination by
these components does not appreciably add ele-
ments of the group Ta to Vb, or indeed Sr. Ba and
Th are the most enriched elements in both cases.
Contamination effects will obviously be easier to
detect in basalts with originally flat MORB-
normalized trace element patterns. For intra-plate
basalts (Fig. 6.37) with originally 'humped' shaped
patterns, crustal contamination effects would be
very much more difficult to discern. Thus for

6
5
4
3
•
2
a
6
4
•
2
•
•
• ••
••
•
•
• ...
• • •
..., .•
• • •.1' •
,..--.
..,.- ..
•
•••
•
•
•
•
• •
•
K20
•
•
•..•
•
•
•
..- .
•• •, • MgO
.....
• ••••••
.. ., .
· ..·rr·.... •
..... .
•
•
•
..
..... It .1P•••••
•• •••
55
• •
•
•
•
Figure 7.22 Variation of % K20 and % MgO versus Si02 for
volcanic rocks form the Calipuy group of Peru (data from
Pitcher et al. 1985).
Andean magmas, generated from enriched sub-
continental lithosphere sources, quantification of
the role of high-level crustal contamination using
trace element geochemistry alone may prove to be a
near-impossible task.
In studying suites of subduction-related basalts,
bivariate diagrams based on trace element ratios
have been found to be useful in separating
subduction-related from mantle components in the
petrogenesis of the magmas (pearce 1982). Figure
7.26 is such a diagram showing the variation of
Th/Yb versus Ta/Yb (Pearce 1982, 1983). Yb is the
denominator in both ofthese ratios, and this has the
effect oflargely eliminating variations due to partial
melting and fractional crystallization processes,
allowing attention to be focused on source composi-
tion as a major petrogenetic variable. Mid-ocean
ridge basalts (MORB) and uncontaminated intra-
plate basalts plot within a well defined band with a
slope of unity, as mantle enrichment events appear
to concentrate Ta and Th equally. In contrast,
island-arc and active continental margin basalts are
displaced to higher Th/Yb ratios, presumably
reflecting the influence of subduction-zone fluids
enriched in Th in their petrogenesis. The fact that
active continental margin basalts plot above the
enriched end of the oceanic mantle array in Figure
7.26 would seem to provide strong support for the
involvement of subcontinental lithosphere in their
petrogenesis. Contaminated continental intra-plate
basalts also plot in a similar position, however, and
thus once more we are faced with the problem of
distinguishing between the effects of subduction-
zone fluids and those of near-surface crustal con-
tamination in producing the observed trace element
characteristics of the magmas.
Tables 7.2-5 include trace element data for
volcanic and plutonic rocks from the Andes. In
general, the active continental margin magmas
appear to show greater degrees of enrichment of a
whole range of incompatible trace elements com-
pared to oceanic island-arc basalts, which may
reflect the combined effects of derivation from an
enriched mantle source and crustal contamination.
7.7.s Radiogenic isotopes
Nd-Srisotopes
Isotopic compositions of Sr, Nd and Pb provide
some of the most useful information for elucidating
magmatic processes at convergent plate boundaries,
because the various source components involved
have such contrasting isotopic signatures. Figure
7.27 shows the variation of 143Nd/144Nd versus
87Sr"s6Sr for volcanic rocks from the northern,
central and southern volcanic zones of the Andes
(Hawkesworth et al. 1982, James 1982, Thorpe et
al. 1984, Hickey et al. 1986) compared with fields
for MORB, OIB and oceanic island arcs. Data for
the NVZ in Ecuador and Colombia and for the SVZ
are displaced to the low 143Nd/144Nd side of the
MORB field, falling within the field of oceanic
island basalts (OIB). Clearly, these .data cannot be

Table 7.2 Average major and trace element compositions of Andean volcanic rocks. compared with those from the island arcs
of the south-west Pacific (SWP) (data from Ewart 1982).
Basalt Basaltic andesite Andesite
%
Si02
Ti02
AI20 3
Fe203
FeO
MnO
MgO
Cao
Na20
K20
P20 5
ppm
Rb
Ba
Sr
Zr
La
Ce
Y
Yb
Cu
Ni
Co
Cr
V
Nb
Pb
Hf
Andes
51.05
1.14
18.57
3.42
5.48
0.16
5.54
8.87
3.98
1.42
0.38
49.9
345
608
162
16.3
41.6
31.0
2.29
30.0
57.9
29.6
67.9
187
2.9
SWP
50.07
0.85
16.23
3.23
6.75
0.18
7.84
10.82
2.51
1.24
0.28
29.1
364
628
69.7
11.6
25.9
19.7
1.54
121
104
43.0
273
300
5.3
7.2
1.3
Andes
53.90
1.27
17.50
3.13
5.39
0.15
5.35
7.68
3.67
1.62
0.35
45.4
676
644
179
24.6
51.3
25.4
2.32
49.6
67.4
30.5
202
220
12.5
3.67
accounted for simply by partial melting of a
depleted asthenospheric mantle wedge (MORB
source mantle) enriched in radiogenic Sr by slab-
derived fluids, as is the case for many intra-oceanic
island arcs (Hawkesworth & Powell 1980, Wilson
& Davidson 1984). Instead, petrogenetic models
could involve partial melting of a subduction-
modified enriched mantle source (subcontinental
lithosphere; Pearce 1983) or contamination of
primary magmas derived from subduction-
modified MORB source mantle with a continental
crustal component. However, extensive crustal
contamination of the NVZ and SVZ lavas appears
SWP
54.19
0.8.3
17.07
3.25
5.68
0.16
5.24
9.08
2.92
1.30
0.26
30.3
402
561
105
20.2
36.4
23.3
1.57
105
44.9
29.7
110
235
6.5
8.0
1.75
Andes
59.89
0.95
17.07
3.31
3.00
0.12
3.25
5.67
3.95
2.47
0.31
75.4
886
648
195
38.0
66.8
12.2
1.94
40.0
38.6
18.6
48.4
125
5.46
SWP
59.09
0.73
16.83
2.82
4.16
0.13
3.83
7.05
3.41
1.70
0.23
41.2
479
516
138
25.4
44.0
24.7
1.94
51.8
34.4
21.3
87.4
154
6.3
9.9
2.7
to be ruled out by combined Sr-O isotopic studies
(Section 7.7.6; see also James 1982, Harmon et aI.
1984) and thus their isotopic compositions may give
a good indication of the isotopic characteristics of
the subduction-modified mantle wedge. In con-
trast, the CVZ lavas are characterized by much
more varied isotopic compositions with higher
87Srfl6Sr and lower 143Ndll44Nd. These data un-
equivocally require contamination of mantle-
derived magmas by the continental crust (Hawkes-
worth et al. 1982, James 1982, Harmon et aI. 1984,
Thorpe et aI. 1984). This is consistent with the
observation made in Section 7.2 that the CVZ is

Table 7.3 Major and trace element analyses of basaltic andesites from the northern (NVZ) central (CVZ) and southern (SVZ)
active volcanic zones of the Andes,
NVZ CVZ SVZ
b West a East a a
%
Si02 55,72 55,50 54,22 52,41 54,35 54,88
Ti02 0,89 0.81 0,95 2,02 0,93 1,33
AI20 3 16,89 15,20 16,02 16,25 18.16 16,50
Fe203 8,72 0.90 8,46 9,27 8,50 3,76
FeO 6,00 6,56
MnO 0,10 0,13 0,13 0,14 0,14 0,19
MgO 5,12 8,32 7,66 6,03 5,60 3,28
CaO 7.51 7,56 7,88 6,93 8,46 7,25
Na20 3,86 3,35 3,14 3,95 3,35 4.44
K20 1.14 1.16 1,19 2,50 0,72 0,84
P20 5 0,23 0,17 0.20 0.49 0,17 0,21
H2O 0,10 0,28
CO2
ppm
Cr 515 120 144 96 12
Ni 166 81 82 47 40
Rb 18 23 32 63 18,2 17
Sr 640 495 501 633 557 485
Y 13 17 21 25 15
Zr 11O 79 115 238 80
Nb 6 5 11 34 1.9
Ba 729 367 509 224 265
La 13.4 11,6 15,7 39,3 9,8 10,0
Ce 27.0 23,5 35,0 84,2 24,1 25,2
Nd 16,7 15,27 18,7 40.4 14,5 14,6
Sm 3,9 3,64 4,0 7,7 3,01 2,9
Tb 0.4 0.52 0,6 0,9 0,47 0,7
Yb 1,0 1,37 1,7 2,1 1,59 2,4
Hf 2,7 2,03 3,2 5.8 1,7 1,9
Ta 0,5 0,25 0,5 2,9
Th 2,6 2,63 2,5 6,8 2,0 2,0
Data sources: aThorpe et ai, (1984); bMarriner &Milward (1984); CHickey et ai, (1986),
characterized by much thicker crust with a substan- tholeiites through alkali basalts to leucite basanites,
tial Precambrian basement. have largely escaped contamination by the con-
Also shown for comparison in Figure 7.27 are tinental crust as they show primitive characteristics
Nd-Sr isotopic data for Cenozoic plateau basalts with high MgO contents (6-11%). If this is correct
from Patagonia (Hawkesworth ec aL. 1979). These then their isotopic characteristics may reflect those
have been erupted in an extensional tectonic regime of the subcontinental lithospheric mantle, thus
similar to that ofa marginal basin eCho 8) to the east indirectly supporting the largely uncontaminated
of the Andean Cordillera. It is possible that these nature of the NVZ and SVZ magmas.
magmas, which range In composition from Combined Nd-Sr data are not available for the

Table 7.4 Comparison of the geochemical characteristics of
an alkali basalt erupted to the east of the volcanic front in the
CVZ (Thorpe et al. 1984) and a calc-alkaline basalt erupted in
the SVZ (Hickey et al. 1986).
SVZ calc-alkaline CVZ alkali basalt
basalt
%
Si02 50.30 43.49
Ti02 0.85 2.34
AI20 3 18.88 13.43
Fe203 9.56 13.19
MnO 0.15 0.18
MgO 5.91 9.95
CaO 10.59 12.30
Na20 2.95 3.12
K20 0.44 1.42
P20 5 0.14 0.74
ppm
Sc 32
V 219
Cr 112
Co 34
Ni 50 127
Zn 80
Ga 17
Y 16 27
Zr 59 190
Hf 1.4 4.5
Ta 3.5
Nb 2.0 47
Th 0.9 5.4
Rb 7.7 24
Cs 0.58
Sa 146
Sr 437 871
La 6.09 47.5
Ce 15.3 96.1
Nd 9.3 49.0
Sm 2.36 8.8
Eu 0.92
Tb 0.42 11
Yb 1.60 2.0
Lu 0.26
plutonic rocks of the Coastal Batholith of Peru and
therefore direct comparison of volcanic and pluto-
nic suites is not possible. However, Figure 7.28
shows the available Sr isotopic data (Beckinsale et
ai. 1985) for three segments of the batholith,
Arequipa, Lima and Toquepala. The Lima seg-
~
~
c:
0
.c:
(j
~
(j
0
a:
300
200
100
50
40
30
20
10
• Lcvz
.I
() svz
Rb Th Nb La Sr Sm y
r-~~-,-L'-~r-L,~--------I
Pb Sa u Ce Nd
Figure 7.23 Spiderdiagrams for basaltic andesites from the
northern (NVZ). central (CVZ) and southern (SVZ) active
volcanic zones of the Andes. Data from Thorpe et al. (1984)
and Hickey et at. (1986). Normalization factors from Sun
(1980).
ment is emplaced along the axis ofa marginal basin
(Section 7.5) into country rocks consisting of 'new'
crust; lavas, dykes, sills and basic plutons. It is
probable that all of the magmas which built up this
segment were derived by crystal fractionation of
mantle-derived parental basalts with comparatively
little crustal involvement (Atherton & Sanderson
1985), as evidenced by their low 87Sr/86Sr ratios.
However, it is important to realize that the young
crust into which this segment of the batholith is
emplaced will be isotopically similar to the primary
magmas and thus crustal contamination may in fact
be extensive but not detectable isotopically. In
contrast, the Arequipa and Toquepala segments are
emplaced partly into a craton, the Arequipa massif,
composed of Precambrian gneisses, Upper
Palaeozoic and Mesozoic sediments and 400-440
Ma intrusive igneous rocks. This crustal assemb-
lage provides an array of possible sources of

Table 7.5 Analyses of plutonic rocks from the Lima segment of the Coastal Batholith of Peru (data from Pitcher et a/. 1985).
Gabbro Diorite
%
Si02 49.84 58.65
Ti02 0.94 0.81
AI20 3 24.92 16.84
Fe203 1.27 2.76
FeO 4.03 4.63
MnO 0.13 0.15
MgO 2.65 3.66
CaO 10.58 6.01
Na20 2.73 2.85
K20 0.64 2.16
P20 S 0.12 0.17
ppm
Ba 259 564
Ce 28 37
Co 15 20
Cr 11 16
Hf
La 7 14
Nd 17 19
Ni 7 10
Pb 9 15
Rb 19 70
Sc 26 25
Sr 431 352
Th 1 8
V 163 196
Y 15 25
Zn 48 83
Zr 27 120
radiogenic Sr and it seems probable that crustal
contamination of mantle-derived magmas can ex-
plain the observed variations in 87Sr/86Sr initial.
Pb isotopes
Figure 7.29 shows the vanation of 207PbP04Pb
versus 206Pbl204Pb for volcanic rocks from the
northern, central and southern volcanic zones of
the Andes (a) and for plutonic rocks from the
Coastal Batholith of Peru (b). These data define
broadly linear trends which are quite different from
the trend defmed by oceanic basalts (MORB +
OIB), the Nonhern Hemisphere Reference Line
(NHRL) (Han 1984, see also Ch. 9). Data for both
volcanic and plutonic rocks plot to the high-
Granodiorite Granite
69.04 75.58
0.42 0.22
15.03 13.35
1.37 0.90
1.77 0.43
0.07 0.05
1.21 0.69
2.85 1.41
3.49 3.96
4.07 3.90
0.10 0.03
741 595
38 34
8 3
6 3
18 14
19 15
7 14
14 12
159 144
10 5
237 104
21 16
65 18
22 21
28 24
191 85
207PbP04Pb side of the NHRL, similar to the Dupal
group of oceanic islands and to oceanic island-arc
volcanics. Pb isotopic data for the volcanic and
plutonic rocks show extensive overlap, supporting
the contention that the plutonic rocks do indeed
represent the eroded root zones of former active
volcanoes. Additionally, these data do not appear to
support extensive involvement of depleted asthe-
nospheric mantle, similar to the source of Nazca
plate MORB, in the petrogenesis of the magmas.
Instead, the data arrays appear to define a mixing
line between an enriched mantle component (?the
subcontinenetal lithosphere) and Precambrian
gneissic crustal rocks. Precambrian basement
gneiss with very low 206Pbl204Pb clearly appears to
have been involved in the petrogenesis of the

co
c::
o
~
:>2
u
o
c::
30
20
K
WITHIN·PLATE
COMPONENT
Sa Ta
Key
• CVZ basaltoc andes'te
o CVZ al'ah basalt
Ce Zr Sm
Figure 7.24 MORB-normalized trace element diagram (after
Pearce 1983). showing a typical CVZ basaltic andesite and an
alkali basalt erupted in an intra-plate setting to the east of the
active volcanic zone. For the subduction-related basaltic
andesite the dashed line indicates the within-plate compo-
nent (subcontinental lithosphere). while the shaded area
indicates those elements enriched in the sources by
subduction-zone fluids (data from Thorpe et al. 1984).
Figure 7.26 ThlYb versus TalYb plot to show the difference
between subduction-related basalts and oceanic basalts
derived from depleted sources (MORB) and enriched sources
(OIB). Uncontaminated intracontinental plate basalts should
plot in the enriched mantle source region. Vectors shown
indicate the influence of subduction components (S). within-
plate enrichment (W). crustal contamination (C) and fractional
crystallization (F). Dashed lines separate the boundaries of
the tholeiitic (THl. calc-alkaline (CA) and shoshonitic (5) fields
(after Pearce 1983. Fig. 9).
co
c::
o
~
:>2
u
o
c::
magma :contaminant 4 : 1
Figure 7.25 MORB-normalized trace element diagram to
show the effects of crustal contamination by mixing a basalt
magma (with MORB-normalized concentrations of 0.5) with
50% partial melts of greywacke and diorite crustal rocks
respectively in the proportions 4: 1magma:contaminant (after
Pearce 1983. Fig. 7).
10
1.0
ThlYb
0.10
-------
active continental
margins
0.010;-;
.0n;t------;;0.~,0;:;------:,:':
.O:------..J
TalYb

-0
Z
;j
05128
:c 0.5126
Z
'"
..
0.5124
05122

.
'.
'.
....

0
.
'.
.~
field~
for OIB '.
...
....
Key
• CVZ data
o Patagonlan plateau
basalts
IHawkesworth et al
1979)
OCEANIC
ISLAND
ARCS
Figure 7.27 Plot of 143Nd/144Nd versus 87Sr/86Sr for volcanic
rocks from the northern (NVZ). central (CVZ) and southern
(SVZ) active volcanic zones of the Andes. Data from Hawkes-
worth et al. (1982). James (1982). Thorpe et al. (1984) and
Hickey et aL (1986). Field of oceanic island-arc volcanic rocks
from Figure 6.46 and field of oceanic-island basalts (OIB) from
Figure 9.23.
magmas forming the Arequipa and Toquepala
segments of the Peruvian coastal batholith. This is
in good agreement with the crustal models ofCouch
et al. (1981) and Jones (1981), which show a thick
Precambrian crustal layer in southern Peru beneath
the Arequipa and Toquepala segments, and an
extremely thin one beneath the Lima segment.
It is probable that some of the scatter in the
volcanic and plutonic 207PbP04Pb_206PbP04Pb
data arrays is a consequence of multicomponent
mixing involving a depleted MORB source mantle
component and Pb derived from subducted oceanic
sediments, in addition to the subcontinental litho-
sphere component. For example, the Pb isotopic
composition of the NVZ and SVZ lavas could be
modelled in terms of the introduction of Pb derived
from subducted continentally derived sediments
(via subduction-zone fluids) into a MORB-source
mantle wedge.
Segments of the CVZ between 16-18°S and
21-26°S have very distinctive Pb isotopic composi-
tions which may be related to different crustal
contaminants. This is more clearly revealed in
Figure 7.30, a plot of 87Srfl6Sr versus 206Pbl204Pb.
Data from the two segments define remarkably
good linear trends pointing in the direction of
different crustal contaminants. The 16-18°S data
project towards the isotopic composition of 2000
Ma Precambrian basement gneisses (Charcani
gneiss), whereas the 21-26°S data can be explained
in terms of contamination of mantle-derived mag-
mas by late Precambrian - Palaeozoic metamor-
phic and granitoid intrusive rocks. Both trends
project back to the fields ofNVZ and SVZ magmas,
the isotopic characteristics of which may therefore
indicate those of the subduction-modified mantle
wedge. Again, these data appear to confirm the
involvement of an enriched mantle source, the
subcontinental lithosphere, rather than MORB-
I . . >
crusta contamination
Figure 7.28 Variation of 87Sr/86Sr initial ratio for plutonic
rocks from the Lima. Arequipa and Toquepala segments of
the Coastal Batholith of Peru (data from Beckinsale et al.
1985).

.0
a.
..
0
N
:0
f:
lil
(a) Active volcanic zone
Key
15.70
• NVZ
o CVZ, 21-26°5
15.65
x SVZ
15.60
15.45
(b) Coastal Batholith of Peru
Key
15.70
• Lima segment
D Arequipa segment
15.65
• Toquepala segment
15.60
•
~
() ~
() ~
o 0
o 0 0
.000.-'
O.lil)
"00 J
/ • ,,/ Pacific
x x ~~~, sediments
x 0- • • •
xxx ..
xxx
XX •
X X
• •
........
?" D
I ) Pacific
D.D I ... sediments
.~Ir.••
~ ...
WD •
•
•
15.45,~___-;-;-;;-_ _----;+':--L_ _---;-;~_ _ _-:+;--__----;:!-::-___---:-:~___-d--,---___,..L___-.l
16.5 210
Figure 7.29 207PbP04Pb versus 206PbP04Pb for (a) volcanic and (b) plutonic rocks from the Andes. (a) Data from James (1982),
Harmon et al. (1984) and Hickey et al. (1986); (b) data from Mukasa (1986). Fields for southern Peru Precambrian gneiss, Nazca
plate MORB and Pacific sediments from Harmon et at. (1984); Northern Hemisphere Reference Line (NHRL) from Hart (1984).

Key
I MORB, Galapagos, Iceland, Azores
0.712
x
x CVZ, (except Sao Miguel)
21-26°5 II St Helena, Austral, Cananes, AscenSion
III Kerguelen Tristan, Gough
IV Society, Marquesas, Samoa, Sao Miguel (Azores),
2000MA
0.710
Charcani gneiss x
(0740, 16.95) x
Reunion, Rodriquez, St Paul, Amsterdam, Crozet
x CVZ, 21_26°S
• CVZ, 16-18°S
x 0
x
Vi 0.708
• Xx o Martinique
~
'"
<Xl
CVZ, •
';::
•
II' 16-18°5 •
<Xl
0.706
0.704
0.702,':;-;:-_----:~--~:__-~__=_-__::~--_,_J
17.0 18.0
source mantle, in the petrogenesis of Andean
magmas.
Also shown in Figure 7.30 are Sr-Pb isotopic
data for volcanic rocks from the Lesser Antilles
island arc (Davidson 1986). These data also define a
steep trend, but in this case point to a contaminant
with much higher 206Pb/204Pb than the Andean
crust. Davidson has explained this apparent con-
tinental crustal contamination trend in an oceanic
island-arc tectonic setting in terms of contamination
of mantle-derived magmas by terrigenous sedi-
ments intercalated in the arc crust. However,
White & Dupre (1986) favour a source contamina-
tion model to explain these data (Ch. 6).
7.7.7 Stable isotopes
Oxygen
As considered in Section 6.11.6, the analysis of
oxygen isotopes is a powerful tool for tracing the
involvement of continental crustal materials in
magma genesis because of the large differences in
8180 between crustal rocks and rocks derived from
the mantle (James 1981). Figure 7.31 shows the
variation of 8180 with 206PbP04Pb for volcanic
rocks from the northern, central and southern
Figure 7.30 Plot of 87SrJ86Sr versus 206PbP04Pb for volcanic
rocks from the NVZ, CVZ and SVZ of the Andes, to show
comparison with the fields of oceanic basalts (MORB + OIB)
(White 1985). (Data for Andean volcanic rocks from James
(1982). Harmon et al. (1984) and Hickey et ai, (1986),) Shown
for comparison are isotopic data from Martinique, Lesser
Antilles (Davidson 1986). a suite of contaminated island-arc
magmas,
volcanic zones of the Andes Games 1982, Harmon
et al. 1984). 8180 is lowest in the rocks ofthe SVZ,
ranging from 5.2 to 6.8%0, indistinguishable from
the oxygen isotopic composition of fresh MORB
and OIB (Kyser et al. 1982). It thus seems
reasonable to assume that the SVZ lavas represent
essentially uncontaminated magma compositions,
the isotopic characteristics (Sr, Nd, Pb, 0) of
which reflect those of the subduction-modified
mantle wedge. The NVZ lavas are relatively
homogeneous in terms of 8180 and overlap with the
IOW-&180 end of the field for CVZ rocks from 21 to
26°S, which show evidence of contamination by a
high-&180 crustal component. CVZ rocks from 16
to 18°S define a completely different trend, project-
ing towards a Precambrian gneissic component
with moderate &
180. The different trends of the
16-18°S and 21-26°S segments of the CVZ mirror
those in Figure 7.30, the Sr-Pb isotope diagram,
clearly supporting the involvement of different
crustal contaminants, old crust with high Rb/Sr,
low UlPb and moderate 8180, and young crust with
high Rb/Sr, high UlPb and high 8180 respectively.
The NVZ lavas appear to have undergone slight
contamination by a young crustal component
broadly similar to that involved in the petrogenesis
of the CVZ lavas from 21 to 26°S.

12
11
10
8
6
old crust-
high Rb/Sr. low U/Pb.
moderate !)'BO
young crust - {;>
high Rb/Sr. Ii
high U/Pb.
high !)'BO
@.16--1
!V
8
SUBDUCTION-MODIFIED
MANTLE WEDGE
186 19.0
Figure 7.31 Plot of &'80 versus 206PbP04Pb for volcanic
rocks from the northem (NVZ). central (CVZ) and southem
(SVZ) volcanic zones of the Andes (after Harmon et al. 1984.
Fig. 9. p. 818)
7.8 Detailed petrogenetic model
Most recent studies (Hawkesworth & Powell 1980,
Perfit et al. 1980, Arculus & Johnson 1981, Kay
1984, Wilson & Davidson 1984, Arculus & Powell
1986) have attributed the main features of island-
arc basalt geochemistry to variable contributions
from two main source components; the astheno-
spheric mantle wedge overlying the subducting slab
of oceanic lithosphere and a metasomatic compo-
nent, either a hydrous fluid or a partial melt,
derived from the subducted oceanic crust. In the
active continental margin tectonic setting two
additional components are involved, the crust and
mantle portions of the continental lithosphere,
making this one of the most complex magma
generation environments on Earth.
It is generally accepted that the continental
lithosphere is thicker than the oceanic lithosphere,
and in Andean-type margins it is probable that
much if not all of the mantle wedge overlying the
subducted slab has lithospheric characteristics.
This subcontinental mantle wedge may have very
DETAILED PETROGENETIC MODEL 223
different trace element and isotope geochemical
characteristics from the underlying depleted asthe-
nosphere, particularly if it has formed part of a
stable continental keel for several billion years.
Specifically, it may be heterogeneously trace ele-
ment enriched due to the migration of partial melts
generated during previous intra-plate magmatic
events (Ch. 3). Addition of slab-derived fluids to
such enriched mantle will induce partial melting if
solidus temperatures are exceeded, and the resul-
tant magmas should have distinctive trace element
geochemistries (Section 7.7.3).
Individual subduction systems differ in signifi-
cant ways, and therefore it is unrealistic to expect
any simple general model to explain all the charac-
teristics of all arcs, both oceanic and continental.
For example, there may be significant pre-
subduction heterogeneity in the mantle wedge and
the geochemistry of slab-derived fluids may vary as
a consequence of variable degrees of submarine
alteration of the oceanic crustal layer, and varia-
tions in the proportions and geochemical character-
istics of any sediments that may be subducted.
Nevertheless, there is one characteristic feature
which appears to be common to all instances of
subduction-related magmatism - the transfer of
Sr, K, Rb, Ba, Th ± Ce, P and Sm to the mantle
wedge by partial melt or fluid-transfer processes
associated with the dehydration of the subducted
slab (Anderson et al. 1980, Hawkesworth & Powell
1980, Wilson & Davidson 1984). This provides the
critical link between the physical process of sub-
duction and arc magmatism.
Once primary magmas have been generated by
partial melting of the subduction-modified mantle
wedge, they must subsequently rise through a thick
section of continental crustal rocks, up to 70 km in
the case of the CVZ of the Andes. Crustal
contamination seems inevitable and the subsequent
geochemical evolution of the magmas must be
dominated by assimilation - fractional crystallisa-
tion processes (AFC) (DePaolo 1981). Thus active
continental margin magmas should in general have
distinctive Sr, Nd, Pb and 0 isotopic signatures,
reflecting the nature of the specific crustal compo-
nent with which they have interacted. This may be
upper or lower crust, young crust or ancient

Precambrian crust, each of which will have differ-
ent isotopic characteristics. Where magmas rise
through young crust Sr, Nd and Pb isotopic data
may give the misleading impression that the mag-
mas are uncontaminated. This is because young
crustal rocks can have isotopic characteristics quite
close to those of the mantle-derived magmas,
particularly so if they represent island-arc sequ-
ences newly accreted to the continental margin. In
such a situation, while AFe processes may have
operated, the isotopic composition of the magmas is
not modified significantly.
In addition to the effects of high-level crustal
contamination, subduction-related magmas may
also inherit an isotopic signature from the continen-
tal crust via the subduction of terrigenous sedi-
ments. This has been clearly demonstrated for
some intra-oceanic island-arc magmas (Section
6.11), the isotopic compositions of which could not
have been modified by direct interaction with the
continental crust. In an Andean tectonic setting it is
probable that contamination in both of these
environments contributes towards the overall con-
tinental crustal fingerprint, but isotopic and trace
element data do not allow us to separate these
effects.
Primitive basaltic magmas generated in the
mantle wedge rise, because they are less dense, to
depths at which there is a zero density contrast
between the magma and the wall rock. In oceanic
island arcs this may be only a few kilometres from
the surface, whereas in a continental margin
environment it is most likely to be in the deep crust
close to the Moho (crust/mantle boundary). The
continental crust, by virtue ofits lower density than
the oceanic crust, thus acts as a filter causing the
subduction-zone magmas to stagnate, become con-
taminated and fractionate at much deeper levels.
The comparative rarity of basaltic lavas in continen-
tal margin arcs may thus reflect their inability to
rise through the continental crust, rather than a
lack of basaltic primary magmas.
Young immature intra-oceanic island arcs are
characterized by relatively high proportions of
tholeiitic mafic volcanic rocks, the trace element
and isotopic compositions of which reflect deriva-
tion from depleted asthenospheric mantle with
minor additions of slab-derived material (Perfit et
al. 1980, Arculus & Johnson 1981, Gill 1981). In
contrast, more mature island arcs and continental
margin arcs underlain by thicker crust erupt greater
proportions of more silicic volcanic rocks. Addi-
tionally, in these arcs, although tholeiitic, calc-
alkaline and shoshonitic series volcanics are all
represented, calc-alkaline and shoshonitic types
predominate. This may reflect the combined effects
of more enriched mantle sources and crustal
contamination in the petrogenesis of the magmas.
As an oceanic island arc evolves with time,
repeated influx of magma causes the crust to
thicken and thus the depth of stagnation of
primitive basaltic magmas to increase (Leeman
1983). Thus, in some instances, the changeover
from dominantly tholeiitic to calc-alkaline arc
magmatism may not necessarily reflect any funda-
mental differences in the primary magma chemis-
try, but simply differences in fractionation condi-
tions. For example, the evolution of basaltic
magmas fractionating at shallow depths outside the
stability field of amphibole will be dominated by
anhydrous assemblages involving plagioclase, oli-
vine, orthopyroxene, clinopyroxene and magnetite,
and the magmas may consequently evolve along a
tholeiitic liquid line ofdescent. However, at greater
depths crystal fractionation of hydrous basic mag-
mas will be dominated by amphibole, which has
been postulated by many authors to be fundamental
in producing calc-alkaline magma chemistries
(Eggler & Burnham 1973, Cawthorn & O'Hara
1976, Allen & Boettcher 1978). However, Hawkes-
worth & Powell (1980) suggested, for the Lesser
Antilles island are, that tholeiitic and calc-alkaline
magmatism was triggered by the release of hydrous
fluids and partial melts respectively from the
subducted slab, and thus that the parental magma
compositions of the two series do differ. This
remains a matter for further detailed study.
In a region of particularly long-lived arc magma-
tism the thermal effects of basaltic magma influx
into the base of the crust become important
(Patchett 1980) and may eventually cause crustal
anatexis (partial melting). As considered in Section
7.4, partial melting of lower crustal gneisses could
produce silicic magmas, and many authors have

attributed the ignimbrite eruptions of the central
Andes to such a mechanism. Additionally, mantle-
derived magmas may mix with such crustal melts
while simultaneously undergoing crystal fractiona-
tion (DePaolo 1981).
The Andean active continental margin has pro-
vided a particularly useful natural laboratory in
which to study the interaction between subduction-
related magmas and the continental crust, as it
shows marked variations from north to south in the
subduction-zone geometry, the volume and prove-
nance of subducted sediments and in the thickness,
age and composition of the overriding continental
crust. The uniformity of the volcanic front relative
to the position of the Peru-Chile trench along the
whole length of the Andean Cordillera implies an
intimate association between volcanism and sub-
duction, which is most easily attributed to the role
of slab-derived fluids. Chemical and isotopic data
for the most primitive basalts erupted in all three
active volcanic zones (NVZ, CVZ and SVZ)
suggest that magma genesis is initiated as slab-
derived fluids, enriched to variable extents in
incompatible elements of low ionic potential,
radiogenic Sr derived from sea water and radio-
genic Pb derived from subducted· sediments, in-
vade the mantle wedge. Partial melting of this
enriched peridotite source region to different de-
grtes produces the primitive mafic magmas which
are parental to the range of rock types observed in
all these volcanic zones. The Sr, Nd, Pb and 0
isotopic characteristics of the NVZ and SVZ lavas
have been interpreted in terms of derivation from a
subduction-modified enriched mantle source with
very little crustal contamination (Thorpe et al.
1981, 1984; Harmon et al. 1984; Deruelle et al.
1983). In contrast, the CVZ lavas are more evolved,
with higher 87Sr/86Sr and al80 and lower 143NdI
144Nd ratios, indicating contamination by the
continental crust. There appears to be a particularly
good correlation between the chemistry of the
Andean volcanic rocks and the crustal thickness
and age (Table 7.1). Thus CVZ lavas have the most
obvious continental crustal fingerprint as these
have risen through the greatest thickness of Pre-
cambrian basement gneisses.
FURTHER READING 225
active continental
margin magmas
---AI ~
upper crust
lower crust
.0IIII
~
primary magmas
J l
enriched
subcontinental
lithosphere
depleted asthenosphere
",0-"7"rt;"m,,,
sediment + sea water
basalt + sea water
Continental
crust
Metasomatized
mantle wedge
Slab-derivec!
fluids
Subducted
oceanic crust
Figure 7.32 Flow diagram to summarize the source compo-
nents involved in the petrogenesis of active continental
margin magmas.
Figure 7.32 summarizes the processes and source
components involved in the petrogenesis of active
continental margin magmas.
Further reading
Gill, J.B. 1981. Orogenic andesites and plate tec-
tonics. Berlin-Heidelberg; Springer-Verlag, 390
pp.
Moorbath, S. & R.N Thompson (eds) 1984. The
relative contributions of mantle, oceanic crust
and continental crust to magma genesis. Phil
Trans R. Soc. Lond. A310, 437-780.
Pitcher, W.S., M.P. Atherton, E.J. Cobbing &
R.D. Beckinsale (eds) 1985. Magmatism at a
plate edge. Glasgow: Blackie 328 pp.
Thorpe, R.S. (ed.) 1982. Andesites: orogenic ande-
sites and related rocks. Chichester: Wiley 724 pp.

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