Myron G. Best. Igneous and metamorphic 2003 Blackwell Science

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The intimate association of sheared serpentinite

bodies with mélange (Figure 18.28a) and the common

occurrence of serpentine in the sheared matrix suggest

that serpentinite might play a role in the origin and em-

placement of mélange. Several possibilities exist where

the low strength of serpentinite and its considerable

buoyancy (density 2.5–2.6 g/cm

at atmospheric con-

ditions) in the presence of denser maﬁc–ultramaﬁc

rocks (3.0–3.3 g/cm

) can potentially lead to mobilized

upwelling masses. At transform zones cutting across

oceanic rifts, fracture-enhanced hydration of ultramaﬁc

rocks apparently creates protrusions onto the seaﬂoor

that subsequently break up into serpentinite detritus.

Dredge hauls reveal larger clasts are cemented by car-

bonate to form ophicalcite, whereas ﬁner detritus and

admixed minor amounts of fossiliferous sediment form

“sedimentary” serpentinite. Coherent blocks of the

disrupted ophiolite sequence tectonically mix into this

soft serpentinite matrix. This serpentinite mélange can

subsequently be accreted onto the overriding land

plate where additional internal brecciation and mixing

can occur. Where subducting oceanic slabs liberate

water, peridotite in the overlying mantle wedge can be

hydrated into serpentinite and mobilized into buoyant

diapirs (Figure 11.16; see also Guillot et al., 2000). As

they ascend, possibly along the wedge–slab interface,

fragments of wallrock can be incorporated (Guillot

et al., 2000). Such diapirs in the Mariana–Izu–Bonin

forearc of the western Paciﬁc have emerged as sea-

mounts or “mud volcanoes” as much as 30 km in dia-

meter and 2 km high on the ocean ﬂoor. Recovered

samples from these ediﬁces consist of serpentinite frag-

ments and exotic blocks of rock containing minerals

(lawsonite, crossite, Na-clinopyroxene) equilibrated at

conditions transitional to that of the blueschist facies

(about 400°C, 15–20 km depth) in a mud matrix of

serpentine, chorite, and clay and carbonate minerals

(Fryer et al., 1999; O’Hanley, 1996).

Serpentinite buoyancy likely enhances ascent of

mélange blocks in the Cloos (1982) wedge-ﬂow model

and may be a factor in rapid exhumation of blueschist-

and eclogite-facies rocks equilibrated at depths of

40 km in subduction zones (see below). However,

the exact reason for rapid exhumation that precludes

equilibration to higher T assemblages remains elusive

(Bebout et al., 1996).

18.6.4 Subgreenschist Facies Rocks

Until the pioneering research of D. S. Coombs (1961)

on the thick sequence of volcanogenic clastic rocks in

southern New Zealand (Figure 18.30), little was known

and appreciated regarding low-grade, subgreenschist

Metamorphism at Convergent Plate Margins: P–T–t Paths, Facies, and Zones

597

Calc-alkaline magmatic rocks

and high-T metamorphic rocks

Ultramaﬁc rocks (black) bordered

by lawsonite  albite  chlorite

 blue amphibole assemblages

(hachures)

Prehnite–pumpellyite facies and

unmetamorphosed sediments

Zeolite facies

Christchurch

Greenschist facies

Amphibolite facies

0 100mi

160km

Dunedin

18.30 Regional metamorphic zones and right-lateral Alpine fault, South Island of New Zealand. Redrawn from Landis and Bishop (1972).

facies metamorphism. Although much remains to be

learned, his conception of the zeolite and prehnite–

pumpellyite facies has withstood the test of innumer-

able subsequent investigations of arc rocks all over the

world.

Investigation of these rocks presents several chal-

lenges. Relict fabrics and minerals of the protoliths are

preserved, seriously hampering the interpretation of

overprinting metamorphic mineral assemblages. Newly

equilibrated low-T minerals are typically very ﬁne

grained and commonly occur as fringes on pre-existing

primary mineral grains. Whether new minerals con-

stitute an equilibrium assemblage can always be ques-

tioned and whether they develop at all depends on the

availability of a penetrating ﬂuid; this is particularly

acute in the case of maﬁc magmatic rock protoliths that

are initially volatile free. Whereas more permeable sedi-

mentary and tuffaceous rocks, especially where glassy

(Figure 14.2), can experience more pervasive recrystal-

lization resulting from invaded ﬂuids, interlayered

basaltic and andesitic rocks are less susceptible to

equilibration under changing metamorphic conditions.

This is particularly the case where the rock mass re-

mains undeformed without avenues for invasion of

ﬂuids. The composition of the penetrating ﬂuid can

also be crucial; Thompson (1971) showed that only

minute concentrations of CO

(X  0.0075) in a

aqueous ﬂuid are sufﬁcient to stabilize calcite  kaoli-

nite  quartz in lieu of laumontite. Thus, in many

metamorphic terranes possibly ﬂooded by ﬂuids

sufﬁciently rich in CO

, zeolite-facies and possibly

prehnite–pumpellyite-facies assemblages are absent

and the onset of metamorphism is marked by greenschist-

facies assemblages.

Understanding the paragenesis of the zeolite and

prehnite–pumpellyite facies formed under the lowest

P–T conditions of metamorphism provides a perspect-

ive on transitions into the greenschist and blueschist

facies formed at higher T and P (Figure 14.33) that

occur in nearby geologic settings (Figure 14.24). How-

ever, a lack of well calibrated mineral thermobarome-

ters hinders the interpretation of the lower T rocks and

creates inaccuracies in P–T paths. Equilibration, partial

or complete, at temperatures less than closure tem-

peratures also hinders determination of chronologic

aspects of metamorphism.

Zeolite Facies. Although Eskola recognized the low-T

existence of zeolitic mineral assemblages as early as

the 1930s, it was not until Coombs et al. (1959) docu-

mented regular mineralogical zones in regionally ex-

tensive zeolite-bearing rocks in southern New Zealand

(Figure 18.30) that the zeolite facies was proposed.

Here, during the Triassic, locally as much as 9 km of

volcanogenic clastic material accumulated in a subsid-

ing trough adjacent to a volcanic arc. As the thickening

pile was subjected to burial metamorphism, the reac-

tive, commonly glassy volcanic detritus recrystallized

into very ﬁne-grained zeolites (Figure 14.3) and other

low-T phases, including albite and epidote. The zeo-

lites stilbite and mordenite developed at shallower

depths and lower temperatures, whereas heulandite

and then laumontite and wairakite formed at higher

temperatures. Miyashiro (1994) considers any equili-

brated assemblage of zeolite  quartz to belong to the

zeolite facies.

Prehnite–Pumpellyite Facies. In the deeper part of the

thick Mesozoic pile of arc-derived clastic rocks in

southernmost New Zealand, prehnite and pumpellyite

appear together with laumontite and other zeolites. But

in similar volcanogenic rocks elsewhere in the South

Island of New Zealand, in Japan, and in other areas of

the world, zeolites and quartz are not stable with

prehnite  pumpellyite. (Pumpellyite (Table 14.1 and

Appendix A) is a bluish-green epidote-like mineral

with a composition roughly equivalent to epidote 

actinolite  water.) However, prehnite and pumpel-

lyite occur with epidote  chlorite and other phases

typical of the lower greenschist facies, including acti-

nolite, albite, stilpnomelane, white mica, titanite, and

calcite, or with lawsonite  albite and other phases

typical of the blueschist facies. Coombs (1961) desig-

nated widespread assemblages of minerals that are

transitional between the zeolite and the greenschist or

blueschist facies the prehnite–pumpellyite facies. Maﬁc

rocks of this facies typically have a distinct bluish-green

color because of abundant pumpellyite. Schmidt

and Robinson (1997) have made a detailed analysis

of the transition from smectites to chlorites in meta-

basites equilibrated under zeolite- to greenschist-facies

conditions.

Prehnite is not stable in the highest grade assem-

blage of the prehnite–pumpellyite facies that has coex-

isting actinolite  pumpellyite  chlorite  epidote.

This four-phase assemblage would seem to be imposs-

ible in the three-component ACF system of Figure

18.31, but additional components such as Fe

3

and Ti

may account for its stabilization.

The stability ﬁeld of prehnite  pumpellyite ranges

from less than 200°C to about 300°C and from about

4 kbars to 1 kbar, where the two minerals are stabil-

ized in seaﬂoor metamorphic rocks. With increasing

T, these two phases react with chlorite and quartz to

produce epidote  actinolite, marking the transition

into the greenschist facies. In rocks poor in chlorite

or quartz, pumpellyite or prehnite can persist stably

into the greenschist-facies regime to temperatures

approaching 400°C. Interpretation of maﬁc mineral

assemblages that contain pumpellyite, prehnite, epi-

dote, and actinolite is fraught with problems, probably

in part kinetic (Beiersdorfer and Day, 1995).

598 Igneous and Metamorphic Petrology

18.6.5 Metabasites at High Pressures

In contrast to the generally subtle prograde mineralo-

gical changes in metabasites at intermediate pressures

in Barrovian terranes, more pronounced changes occur

in basaltic protoliths as they are subjected to increasing

P at low to moderate T from the prehnite–pumpellyite

facies through the blueschist and eclogite facies (Figure

14.33). Diverse chemical rock groups are represented

in these high-P facies but maﬁc rocks provide the most

diagnostic mineral assemblages, hence the following

focus on them.

Blueschist Facies. The most obvious mineralogical

contrast between the blueschist and lower P–higher T

facies lies in the nature of the amphiboles in meta-

basites. Whereas green calcic (Ca–Mg–Fe

2

) actinolite

typiﬁes the prehnite–pumpellyite and lower greenschist

facies, the characteristic amphibole of the blueschist fa-

cies is blue sodic (Na–Mg–Fe–Al) glaucophane. This

amphibole has a striking pleochroism in thin section

ranging from colorless to yellow to blue and purple;

in hand sample it is blue to blue-black, hence the

blueschist facies. However, glaucophane solid solu-

tions are stable over a broad range of P–T conditions

not restricted to just those appropriate to the blueschist

facies. Under some conditions, a miscibility gap exists

between actinolite and glaucophane so that they coex-

ist stably in the same rock. Because glaucophane alone

is not indicative of the blueschist facies, its deﬁnition

depends on the stable coexistence of lawsonite plus

glaucophane, in contrast to the coexistence of the com-

positionally equivalent albite  chlorite  actinolite

that is diagnostic of the greenschist facies. (In the long

and controversial history of the validity of the blue-

schist facies, it was known as the more speciﬁc, but

more cumbersome, glaucophane–lawsonite facies.)

For a hydrous mineral, lawsonite is surprisingly dense

(3.1 g/cm

); its composition, CaAl

(Si

)(OH)

·H

is equivalent to anorthite (density 2.76 g/cm

)  water.

In basaltic protoliths, lawsonite has been observed to

crystallize directly from magmatic plagioclase accord-

ing to

18.24 CaAl

·NaAlSi

 2H

O →

plagioclase

CaAl

(OH)

·H

O  NaAlSi

lawsonite albite

Lawsonite appears to be stable to pressures as low as

3 kbar in some assemblages.

Additional diagnostic minerals of higher P–lower T

assemblages in the blueschist facies include aragonite

and jadeitic clinopyroxene. Pure jadeite  quartz is a

stable mineral pair in lieu of pure albite in the higher P

part of the facies. Solid solution with aegerine, diop-

side, and hedenbergite end-members lowers the stabil-

ity P of jadeite and where quartz is absent the stability

P is still less (Figures 16.4 and 16.10). Because of

relatively low temperatures, biotite does not occur,

although the nearly identical (in terms of optical prop-

erties in thin section) stilpnomelane is widespread. The

absence of andalusite and sillimanite in rare blueschist-

facies pelites reﬂects their high-P stabilization. De-

pending on bulk rock composition, chlorite, titanite,

pumpellyite, Ca–Mg–Fe–Mn garnet, and phengitic

white mica are also found in blueschist-facies metaba-

sites (Figure 18.32).

Some mineral assemblages transitional between the

blueschist and other facies depend on the oxidation

state of the rock. Stabilities of amphiboles and epidotes

are especially affected, such as in the reaction (Evans

and Brown, 1987)

18.25 Na-amphibole  epidote 

albite  chlorite  actinolite  Fe-oxide

In this reaction, Fe

3

does not substitute appreciably

in any of the silicate phases in the greenschist assem-

blage on the right and, hence, does not perturb its

stability ﬁeld. However, considerable substitution can

take place in the two phases on the left, extending their

stability ﬁeld to lower pressures than for the Fe

3

-poor

assemblage of glaucophane  zoisite (Yardley, 1989,

p. 106). This effect may explain the intimate interlayer-

ing in some places of schists that are of blue and green

color. Both must have been metamorphosed under the

same P–T conditions, but the schists that contain blue

crossite (a Na–Fe

3

-rich amphibole compositionally

intermediate between glaucophane and riebeckite;

Metamorphism at Convergent Plate Margins: P–T–t Paths, Facies, and Zones

599

Actinolite

Prehnite

Epidote

Pum-

pellyite

 Albite

Chlorite

18.31 ACF diagram for higher T part of the prehnite–pumpellyite

facies. Dark shaded ellipses represent compositions of basaltic

rocks (b) and arc andesites and dacites (a).

Appendix A, no. 17) developed in oxidized layers and

the green actinolite schists in reduced layers that con-

tain little Fe

3

. Widespread terranes of blue crossite-

bearing rocks, including the Sanbagawa metamorphic

belt in Japan (J. Metam. Geol., 1990, vol. 8, no. 4)

and the Shuksan terrane of the Cascade Mountains

of western Washington (Brown, 1986), are believed to

be transitional in terms of P of formation between

Barrovian-type greenschist-facies rocks and blueschist-

facies rocks that contain the diagnostic assemblage of

glaucophane  lawsonite.

Regional terranes of blueschist-facies rocks, some

with mineral assemblages transitional into the lower P

prehnite–pumpellyite facies and/or into the higher P–T

eclogite facies, occur in the main island of New Cale-

donia in the western Paciﬁc Ocean (Yokoyama et al.,

1986), parts of the Alps, and many other subduction

settings (Evans and Brown, 1986). The Franciscan

terrane is probably the best known where blueschist-

facies rocks occur as blocks with less common eclogite

in a typically poorly exposed mud-matrix mélange and

as variably deformed but coherent kilometer-scale

sheets of metasedimentary rock.

Eclogite Facies. Eskola’s studies of the dense eclogites

in the Western Gneiss region of Norway in the early

1900s prompted him to propose a high-P eclogite

facies. However, several problems had to be overcome

before its acceptance today as a valid facies distinct

from other facies. Initially, only metabasites made

essentially of the assemblage omphacite  magnesian

garnet were known (Figure 16.9b). Some petrologists

believed that low water fugacities were responsible for

destabilizing hydrous phases diagnostic of other facies,

such as glaucophane and epidote in the blueschist

facies. Whether glaucophane, white micas, epidote

minerals, chlorite, and other hydrous minerals that

occur in some eclogitic rocks belong with the eclogite

facies assemblage or were stabilized during subsequent,

common retrogressive metamorphism (Figure 14.27b)

was problematic; if these hydrous minerals are primary

they indicate a signiﬁcant amount of water in the

eclogitic system.

Although much has yet to be learned about the

rare but intriguing mineral assemblages of the eclogite

facies, there is consensus that they equilibrate over a

wide range of T, from the 1000°C of mantle-derived

xenoliths to the 500°C of convergent plate associations

—a wider T range than all other facies (Section 15.2.4

and Figure 14.33). Considering the wide range of the

now available chemical rock types (Carswell, 1990,

p. 4), mineral assemblages and their equilibrium P–T

conditions indicate a broad transition between the

eclogite and adjacent facies.

18.6.6 P T t Paths and Tectonic Evolution of

High P/T Terranes

Rocks of the prehnite–pumpellyite, blueschist, and

locally eclogite facies in accretionary wedges record

rapid emplacement of rocks, burial, tectonism, meta-

morphism, and exhumation during plate convergence.

Depositional ages of sediment inferred from rarely

preserved fossils are commonly not much greater than

the isotopic age of metamorphic recrystallization and

thus not much older than the time of exhumation.

Most extensive terranes are Cenozoic or late Mesozoic,

apparently because older ones have been overprinted

by continuing or subsequent higher T metamorphism,

commonly under amphibolite-facies conditions. Such

overprints with only sparse relicts of blueschist-facies

600 Igneous and Metamorphic Petrology

Phengite

0 0.3mm

Chlorite

Titanite

Lawsonite

Glaucophane

Aragonite

Garnet

18.32 Thinly interlayered calc–silicate and maﬁc schists metamorphosed under blueschist-facies conditions, Franciscan Complex, Laytonville

Quarry, Mendocino County, California. Aragonite has variable relief.

assemblages have been noted in several terranes, such

as in the Appalachian Mountains of Vermont (Laird

and Albee, 1981).

The contrasting tectonothermal processes in con-

vergent plate terranes that, on the one hand, preserve

high-P/T rocks and, on the other hand, produce higher

T metamorphism were ﬁrst recognized in pioneering

work by Ernst (see 1988). Most such metamorphic

terranes follow P–T paths that move into greenschist-

and locally amphibolite-facies conditions on the retro-

grade segment of their path (Figure 18.33) because of

signiﬁcant heating before decompression; subduction

of heat-producing continental rocks, as in the Alpine

model in Figure 18.2a, might have been responsible.

In contrast, fewer terranes, such as the Franciscan,

remained refrigerated during exhumation, perhaps

because they were incorporated into rapidly rising, very

buoyant serpentinite diapirs and mélanges spawned in

the hydrating mantle wedge overlying the descending

slab. The corresponding steeper P–T path penetrates

well into the stability ﬁeld of aragonite, while the ret-

rograde path essentially retraces the prograde at only

slightly higher T, avoids greenschist-facies conditions,

and returns into the ﬁeld of calcite at low T, perhaps on

the order of 100°C. This sort of refrigerated path may

account for widespread preservation of metastable

aragonite in Franciscan rocks of appropriate com-

position (Figure 18.32a) and its absence in rocks meta-

morphosed to higher T blueschist-facies conditions.

Although some of the latter never enter the stability

ﬁeld of aragonite, others at lower T do, and because

their path brings them back into the stability ﬁeld of

calcite at signiﬁcantly higher T, on the order of 400°C,

transformation to the lower P polymorph is kinetically

effective (Carlson and Rosenﬁeld, 1981).

18.7 ULTRAHIGH-P METAMORPHIC

ROCKS

The occurrence of coesite, albeit exceedingly rare, in

metamorphosed continental rocks (Figure 15.18) has

prompted the recognition of a distinct type of ultra-

high-P (UHP) metamorphism. These terranes equili-

brated in the stability ﬁeld of this silica polymorph in

excess of about 29 kbars at 800°C (see Figure 18.35).

Even more rare diamond in such terranes constrains

equilibration pressures to at least 38 kbar (about 140 km

depth) at 800°C—the graphite–diamond transition.

Although the rarity of coesite and diamond, which

stems in part from their sporadic preservation, may

limit the recognition of UHP rocks, other mineral

barometers conﬁrm these extreme pressures.

Study of UHP rocks is not an esoteric petrologic

exercise. The ramiﬁcations of burial of substantial

terranes of continental rock to depths 100 km into

the denser upper mantle and their subsequent exhu-

mation challenge conventional tectonic concepts. For

example, the deepest Moho today is only about 80 km

beneath the Himalaya orogen. Does this mean that the

Himalayan continent–continent collision zone is unlike

past collision orogens where UHP rocks have been

found? Or, are geophysical determinations of the depth

of the thickened Himalayan continental crust ﬂawed in

some way? How is it possible for any tectonic process

to push relatively buoyant continental rock into the

denser upper mantle? How are such once deeply buried

rocks exhumed? What do UHP rocks tell us about our

conventional notions of how continents grow?

When the ﬁrst edition of this textbook was pub-

lished in 1982, UHP rocks were unknown. Following

Christian Chopin’s discovery just two years later

(Chopin, 1984) of coesite in silicic continental rocks of

the western Alps and then a few months later by Smith

(1984) in eclogites in the Western Gneiss region of

Norway, such rocks were subsequently found in China

and in a dozen other orogens worldwide (Hacker and

Liou, 1998). The intriguing petrologic implications

and challenges together with the tantalizing tectonic

paradoxes have spawned over one hundred research

articles, several symposium volumes, and at least four

books, clearly making this one of the hottest (or deep-

est, depending on one’s inclinations) topics in geology

of the late twentieth century!

Metamorphism at Convergent Plate Margins: P–T–t Paths, Facies, and Zones

601

P (kbar)

Depth (km)

0 100 200 300 400 500

T (°C)

Arg

Cal

BLUESCHIST

GREENSCHIST

18.33 Schematic P–T paths for high-P/T rocks. Generalized ﬁelds of

blueschist- and greenschist-facies conditions are indicated.

Aragonite–calcite univariant equilibrium curve from Figure

16.1. Long dashed line depicts a possible path taken by many

Franciscan blueschist-facies rocks. Short dashed line repre-

sents path of blueschist-facies rocks in other terranes that

apparently experience greater heating and consequent equili-

bration of greenschist-facies mineral assemblages. Because

of these higher temperatures, aragonite can transform into

calcite that generally doesn’t happen in the refrigerated

Franciscan rocks because of unfavorable kinetics at low

temperatures. Redrawn from Ernst (1988).

18.7.1 Coesite and Diamond: Diagnostic UHP Minerals

UHP minerals and mineral assemblages are typically

preserved in tectonically emplaced lenses and thin

layers intercalated within lower-P rocks of the amphi-

bolite and greenschist facies. P–T paths can be recon-

structed, sometimes only crudely because of the lack of

pertinent experimental and thermodynamic data, using

disequilibrium mineral assemblages, zoned minerals,

and, especially, mineral inclusions within unreactive

host grains, such as pyropic garnet, jadeitic clinopyro-

xene or omphacite, kyanite, and zircon. These strong

mineral grains served as effective “pressure cells” (Spe-

cial Interest Box 18.2) that preserved the UHP envir-

onment of the inclusion during exhumation of the

rocks. Preserved UHP polymorphs of silica and carbon

provide the most direct barometric insights. Most

coesites have been partially to completely pseudomor-

phed by lower P quartz in distinctive “palisade-textured”

aggregates encased within their host. The larger molar

volume of the replacing quartz commonly ruptured

the host mineral grain during exhumation, creating

surrounding radial expansion cracks (Figure 15.16).

Microdiamonds, generally 0.1 mm in diameter,

also occur as inclusions in garnets in rocks of the

Dabie–Sulu terrane in China and the Kokchetav Com-

plex of northern Kazakhstan. In the Western Gneisses

of Norway (see below), because of their extremely

small abundance, diamonds have only been recovered

by acid digestion of large volumes of rock. Skeptics of

UHP metamorphism have argued that these minute dia-

monds are a contaminant from laboratory processing.

(Rocks are usually cut by diamond impregnated saw

blades and diamond grit is used as an abrasive in pre-

paration of thin sections.) Or, diamonds were formed

at high-stress grain boundaries or by nonequilibrium

factors. But the preservation of diamonds and coesites

as inclusions within some sort of mineral host disproves

both laboratory contamination and high-stress grain

boundary nucleation. Others have argued that diamond

and coesite were somehow stabilized by large nonhy-

drostatic stresses, but this “tectonic overpressure” notion

was discounted decades ago when it was realized how

weak rocks are under metamorphic conditions (Sections

8.2.1 and 17.2.6; see also Stöckhert and Renner, 1998).

Relict mineral assemblages that mostly escaped re-

equilibration at lower P as a result of limited access of

volatile ﬂuids and/or absence of deformation corrobor-

ate ultrahigh pressures inferred from coesite and dia-

mond. For example, the Kokchetav rocks are unique

in containing clinopyroxene grains with exsolution

lamella of K-feldspar; potassic clinopyroxene is stable

only under UHP conditions. Deﬁning assemblages in

classic UHP terranes are considered next for the Dora

Maira Massif of the western Alps and the Dabie–Sulu

terrane of east-central China.

602 Igneous and Metamorphic Petrology

Special Interest Box 18.2 Mineral grains as

“ﬂight recorders” of their geologic history

Everyone has become familiar in recent years with

the “black-box” ﬂight recorder sought in crashed

airplanes that can tell of circumstances leading up

to the crash to aid investigators reconstructing the

event. In a similar way, inclusions embedded in a

mineral grain tell about various aspects of its history.

Garnet porphyroblasts are a useful “ﬂight

recorder” that allow the petrologic investigator to un-

ravel the time-dependent thermobarometric path from

rocks caught up in the orogenic wreckage of plate

collisions (Haugerud and Zen, 1991). The typical long

history of growth of porphyroblasts over millions of

years (see Section 18.2 on the Tauern Window) under

changing metamorphic conditions is recorded in

successive compositional zones (Figure 16.34) that can

potentially yield P–T values using thermobarometric

techniques. “Potentially” is used in the foregoing

sentence because of numerous pitfalls in thermo-

barometry of garnets that arise from complexities in

zoning related to, for example, skeletal growth, dif-

fusional adjustments in compositional zones after

growth, especially at high temperatures, and lack

of original equilibration of Ca concentrations. The

assemblage of mineral inclusions incorporated into

poikiloblastic garnets during growth can be assessed

in petrogenetic grids. In addition to this potential

petrologic information is the opportunity for iso-

topic dating of successive zones of the garnet either

using the Rb–Sr isotope system on the garnet itself

or the U–Pb isotopic system on inclusions of zircon,

rutile, and titanite (see Special Interest Box 19.1).

Amazing discoveries of polymorphs of silica and

carbon stable only under mantle pressures occurring

as inclusions in garnet and other mineral “pressure

vessels” spawned a whole new ﬁeld of ultrahigh-P

metamorphic petrology in the last 15 years of the

twentieth century. Tectonic models had to be revised

to reconcile the deep burial and rapid exhumation

of the sedimentary protoliths that host coesite

and diamond. Incredibly, details of the history of

continental-scale orogens are obtained from micron-

size mineral inclusions!

In a further step, minute mineral inclusions in dia-

monds hosted in peridotite and eclogite xenoliths

in kimberlite (Section 13.12.2) reveal information

about the mineralogical nature of the mantle.

Fluid inclusions in a wide variety of hosts (Spe-

cial Interest Box 16.3) provide useful information

regarding P–T–X

ﬂuid

conditions of mineral growth

and deformation. Melt inclusions provide data on

volatile solubilities in melts at elevated P and T and

their impact on explosive volcanism and other geo-

logic phenomena (Section 4.3).

18.7.2 Dora Maira Massif in the Western Alps

The tectonic collage of metamorphic rocks that makes

up the Dora Maira Massif in the western Alps was

created during late Cretaceous–early Tertiary collision

of continental plates—the European on the north and

one or more continental plates to the south (Figures

18.6 and 18.24). (Massif is from the French massive and

connotes a large block of mountain terrain made of

relatively coarsely crystalline rocks.) UHP rocks occur

as lenses within a stack of tectonically emplaced nappes

and slices of Paleozoic schists, gneisses, marbles,

quartzites, eclogite, and Mesozoic ophiolite (Figure

18.34). Whereas protoliths in circum-Paciﬁc high P/T

terranes are mostly oceanic rocks and near-trench

immature sediments (feldspathic and lithic debris),

Alpine protoliths are continental granitoids and more

differentiated shelf-deposited sedimentary rocks, such

as pelites and quartzites. These rocks appear to have

been metamorphosed to the blueschist and eclogite

facies but a widespread, variable greenschist-facies

overprint took place in the middle Tertiary. The tec-

tonic slices experienced different P–T histories; partly

exhumed UHP and high-P layers appear to have been

thrust over lower-P ones before ﬁnal exhumation of the

entire sequence (Harley and Carswell, 1995).

Inclusions of coesite are hosted within unusually

pure (90–98 mol % end-member) pyrope poikiloblasts

as much as 20 cm in diameter whose dominant inclu-

sion is kyanite. These nearly colorless to pink pyrope

porphyroblasts occur in a kyanite-white mica–talc

quartzite, known locally as white schist. Zircon, rutile,

and jadeitic (70 mol % end-member) clinopyroxene

are accessories. The unusually high Al, Si, and Mg

and virtual absence of Ca and Na in these white schists

indicates an unusual protolith; some sort of sediment-

ary evaporite has been suggested but a metasomatized

granitoid seems more likely. The coesite-bearing white

schists occur as discontinuous layers and boudins cen-

timeters to meters thick hosted within greenschist-facies

Metamorphism at Convergent Plate Margins: P–T–t Paths, Facies, and Zones

603

0km4

0 miles 3

7°20′E

44°35′N

Miscellaneous

metam rocks

Chiefly HP

eclogite

UHP unit; lenses of

coesite-pyrope quartzite

in black

Chiefly blue-

schist facies

Ophiolite and

associated rocks

18.34 Geologic map and cross-section of the southern Dora Maira Massif in extreme northwestern Italy of the western Alps. See location in

Figure 18.6. Each of the map units consists of a variety of rock types but dominated by that indicated. In addition to small lenses of

coesite–pyrope quartzite, the UHP unit contains orthogneiss, marble, a variety of schists, and local eclogite. The slightly generalized cross

section shows the structure of the post-metamorphic dome along a west–east line at approximately 44° 30′N. All contacts appear to be

tectonic. Redrawn from Michard et al. (1995).

gneiss that lies in a sequence of metapelites, marbles,

and metabasites.

P T Path. Schertl et al. (1991) document critical

mineral zoning, inclusions, and reaction textures that

deﬁne the P–T path taken by the coesite-bearing white

schists during their ultradeep burial and subsequent

exhumation. These constraints can be followed in the

petrogenetic grid in Figure 18.35. The initial rock was

likely composed essentially of Mg-chlorite  quartz.

604 Igneous and Metamorphic Petrology

P (kbar)

High-P (HP)

Ultrahigh-P (UHP)

Depth (km)

Normal continental crust Upper mantle

175

150

125

100

100 300 500 700 900 1100

T (°C)

And

Sil

Dora Maira

Si 3.09

Si 3.26

Si 3.55

Tlc



Tlc



Jd  Qtz

Gln  Zo  Qtz

Chl + Qtz

Tr  Ab  Chl

BLUESCHIST

Arg

Cal

Mg-Car

Graphite

Diamond

Tlc  Jd

Gln

rtz

Coesite

Ell

Dabie–

Sulu

ECLOGITE



10°C/km

5°C/km

Kokchetav

18.35 Petrogenetic grid for select univariant reactions pertaining to high-P (HP) and ultrahigh-P (UHP) metamorphism. UHP metamorphism

takes place in the stability ﬁeld of coesite, HP in the quartz ﬁeld. Broad P–T conditions for the eclogite and blueschist facies are indi-

cated. The tight clockwise P–T path for metamorphism in the Dora Maira Massif is indicated by the shaded arrow. Peak metamorphic

conditions for the UHP rocks of the Kokchetav Massif and Dabie–Sulu chieﬂy metasedimentary rocks are shaded. Wedge-shaped areas

ornamented by dots alongside lines are stability ﬁelds of ellenbergerite (Ell) and Mg-Carpholite (Mg-Car). Light lines show reference 5

and 10°C/km geothermal gradients. Dashed lines show Si content of phengitic mica that serves as a barometer (Massone and Schreyer,

1989). Most reaction lines drawn for unit water activity. Redrawn after Schertl et al. (1991), Harley and Carswell (1995), and Zhang et al.

(1997). Some univariant lines are from ﬁgures in this textbook.

The small amount of K in the rock probably stabilized

some additional phengitic white mica while excessive

Al stabilized kyanite. An inclusion of a phengite grain

within a relict kyanite that is in turn hosted within a py-

rope poikiloblast deﬁnes a point at 16 kbar and 560°C

on the prograde P–T path. Because of the P-sensitive

substitution SiMg b AlAl in phengite, the concentra-

tion of Si serves as a useful barometer where it coexists

with another Al-phase, such as kyanite. The phengite

inclusion contains 3.26 Si per 12 oxygen in the formula

unit, indicating equilibration at 16 kbar. At this P, the

T-sensitive reaction chlorite  quartz  talc  kyanite

occurs at about 560°C. Further evidence for this dis-

continuous reaction lies in the absence of any relict

Mg-carpholite (Table 14.1) inclusions in any host

mineral. The next thermobarometric constraint on the

prograde P–T path is provided by the occurrence of

inclusions of purple ellenbergerite, a newly discovered

hydrous Mg–Al–Ti–Zr silicate with extensive solid

solution to a phosphate end-member that occurs as

inclusions in pyrope porphyroblasts. Silica-rich ellen-

bergerite inclusions probably formed by the reaction

chlorite  kyanite  rutile  H

O → ellenbergerite

because relics of the reactant phases are also found as

inclusions around those of ellenbergerite. Its narrow

wedge-shaped stability ﬁeld constrains the prograde

path to about 30 kbar and 700°C after the rock had

passed above the quartz → coesite transition. Peak P–T

conditions in the Dora Maira rocks are preserved in

Si-rich phengite and small pyrope porphyroblasts that

contain abundant coesite inclusions. Because all four

phases in the T-constraining reaction talc  kyanite 

pyrope  coesite appear to be compatible in the rock,

equilibration must have been near 800°C, or less if

water activity was much less than 1. The 3.55 Si per

formula unit in the coexisting phengite constrains P to

about 37 kbar, just short of the diamond stability ﬁeld.

In rare cases, talc  jadeite, instead of the chemically

equivalent glaucophane, occur in mutual contact, sug-

gesting their stable coexistence and conﬁrming the

peak conditions. However, retrograde glaucophane

does occur locally in cracks within garnets. A tight

restriction on the retrograde path is provided by wide-

spread partial replacement of small pyrope porphyro-

blasts by the assemblage phlogopite  kyanite  quartz

 scapolite. Except for scapolite, this assemblage can

be related to the univariant reaction talc  phengite →

phlogopite  kyanite  quartz. Schertl et al. (1991) be-

lieve this reaction was shifted to slightly lower temper-

atures as a result of interactions with hydrous ﬂuids

that contained Na, Ca, and Cl, reducing the water

activity to 1 and stabilizing scapolite. Fine-grained

retrograde white micas, apparently equilibrated with

quartz  kyanite  chlorite, contain 3.09 Si per for-

mula unit, constraining the retrograde path to near

7 kbar at 500°C, appropriate to depths of typical lower

continental crust. Neither andalusite nor sillimanite

occur in the Dora Maira rocks, which means that the

retrograde path, after the rocks arrived at a depth of

about 20 km, must have followed a more normal ther-

mal gradient of about 20°C/km.

There are at least two alternate interpretations for

the origin of the greenschist-facies gneiss (quartz–albite

–alkali feldspar–biotite–white mica–chlorite–epidote)

adjacent to the coesite-bearing rocks. Either the gneisses

were never subjected to UHP metamorphism and were

tectonically intercalated with the coesite-bearing rocks

during their transit through the shallow crust, or the

gneisses were pervasively retrograded from an UHP

condition during exhumation. The latter interpretation

is strengthened by the occurrence of relic phengitic

mica grains that were equilibrated at high P and by rare

armored relics of grossular-rich (85 mol %) garnet and

rutile in almandiferous garnet. Figure 18.35 shows that

grossular-rich garnet is stable with rutile and coesite

near 30 kbar and 800°C.

18.7.3 Dabie–Sulu Terranes

These terranes in east-central China are a segment of

a 2000-km-long continental suture zone deﬁned by

high-P eclogite- and blueschist-facies rocks between

two Precambrian cratons (Figure 18.36). Following

subduction of the seaﬂoor in the late Paleozoic,

the northward moving Yangtze continental mass was

Metamorphism at Convergent Plate Margins: P–T–t Paths, Facies, and Zones

605

18.36 Geologic map of the Dabie–Sulu UHP terrane and associated

rocks in the Triassic suture zone between the Precambrian

Sino-Korean and Yangtze Cratons in east-central China.

Redrawn from Wang et al. (1995). Heavier lines are faults,

including a major north–south strike-slip fault that has dis-

placed the UHP Sulu terrane from the once continuous Dabie

terrane.

112°E 120°E

0 300km

0 200miles

Dabie

CRATON

YANGTZE

Sea

Sulu

36°N

32°

Yellow

Beijing

SINO-

KOREAN

CRATON

Greenschist

facies

Amphibolite-

granulite facies

Blueschist

facies

High-P eclogite

UHP rocks

jammed below the edge of the Sino-Korean continent

in the late Triassic, 240–230 Ma, creating the UHP

rocks now exposed over an area of about 1000 km

Associated tectonites to the north in the Dabie terrane

are amphibolite- and granulite-facies felsic gneisses and

migmatites, whereas to the south, across a possible

detachment surface marked by mélange, more varied

rocks include felsic gneiss, paraschist, marble, ultra-

maﬁc rocks, and lenses and layers of eclogite hosted

within them.

Coesite and coesite relics occur as inclusions in gar-

net, clinopyroxene, kyanite, dolomite, and zircon. Rare

microdiamond inclusions and armored relics of grossular-

rich garnet and rutile in eclogite garnets indicate pres-

sures 30 kbar at 700°C (Figure 18.35). UHP rocks

are pervasively overprinted by amphibolite-facies assem-

blages. Bolin et al. (1999) believe the UHP rocks were

exhumed in three stages between 230 and 120 Ma.

18.7.4 Evolution of UHP Terranes

Until the 1980s, subducting plates were believed to be

essentially restricted to those of oceanic lithosphere;

continental lithosphere with its thick, signiﬁcantly

less dense crust was deemed to be just too buoyant

to subduct into the mantle. In the classic continent–

continent collision zone of the Himalaya, the Eurasian

plate in Tibet is underthrust at a shallow angle by the

Indian continent so that the deepest part of the Moho

lies at depths of “only” 70–80 km (Figure 13.33); the

underthrust continental plate does not penetrate into

the mantle. However, the dogma that continental plates

cannot subduct was overturned in the 1980s with

discoveries of coesite and diamond in continental

protoliths that must have been buried twice as deep, to

depths of 120–160 km (Figure 18.35).

In trying to understand the tectonothermal evolu-

tion of UHP terranes it is signiﬁcant that not only were

they buried to great depth but they are also restricted

to continental, not oceanic, protoliths in elongate

continent–continent suture zones. Protoliths followed

relatively refrigerated P–T paths during burial and

exhumation. After initial burial along an approximate

trajectory of 10°C/km to about 560°C at 55–60 km, the

Dora Maira rocks were subsequently transferred to

800°C at about 130 km (Figure 18.35) along a trajec-

tory of a little more than 3°C/km. Decompression

followed along a similar trajectory to about 30 km.

This path precludes sufﬁcient heating to create partial

melting—hence the virtual absence of coeval silicic

magmatism (Figure 18.6) like the Himalaya granites

(Figure 13.33). More maﬁc calc–alkaline magmatism of

the Andean type (Section 13.7) is also absent because

there is no subducting and water-liberating oceanic

crust to foster generation of magmas in the arc mantle.

What seems to be required is rapid subduction and

return exhumation of cold continental lithosphere.

Just how rapid is revealed by chronologic analyses

of minerals in assemblages that also yield thermobaro-

metric data on P–T paths. SHRIMP analyses (see

Special Interest Box 19.1) for U–Th–Pb isotopes in

titanite grains in the Dora Maira rocks yield an age of

35.1  0.9 Ma for the metamorphism at maximum P

(35 kbar near 125 km depth; Rubatto and Hermann,

2001). Retrograde titanites equilibrated at about 35 km

depth yield an age of 32.9  0.9 Ma, indicating as-

tounding exhumation rates of a few centimeters per

year (the exact rate must take into account the uncer-

tainty errors in the chronologic and barometric ana-

lyses). Subsequent slower exhumation at about 0.5 cm/y

brought the UHP rocks to a depth of about 10 km and

250°C at 29.9  1.4 Ma, based on annealing of ﬁssion

tracks in zircon. Chronologic data on the coesite-bearing

rocks at the western end of the middle Tertiary

Himalaya and diamond-bearing rocks of the 500 Ma

Kokchetav and 210 Ma Dabie–Sulu have larger analyt-

ical uncertainties but are of similar magnitude (centi-

meters per year).

Such astounding rates of exhumation of low-density,

refrigerated continental rocks from mantle depths

appear to be required for preservation of UHP

metamorphism. (One is reminded of the response

of a basketball pushed into the water in a swimming

pool; once released the buoyant ball rapidly pops

back to the surface.) One possible scenario (O’Brien

et al., 2001) involves break-off of the leading, denser

oceanic part of the descending slab that was acted

on by a slab-pull force earlier in the subduction

process. Once freed of its denser leading part, the

buoyant continental material can pop up to the top of

the mantle at rates of centimeters per year. Subsequent

slower rates of exhumation occur as more normal iso-

static rebound takes place in the imbricate stack of

thrust slices.

SUMMARY

Metamorphic zones, facies, and P–T paths provide im-

portant constraints on dynamic tectonic and thermal

phenomena at convergent plate junctures during

orogeny. The nature of the plates, whether continental

or oceanic, their thermal state and thickness, rate of

convergence, distribution of radioactive heat sources,

and rate of unrooﬁng govern the conductive thermal

evolution of orogens and the metamorphic terranes

in them. Faster convergence rates tend to counteract

the heating effect via thermal relaxation of thickened

crust that largely controls the prograde path of meta-

morphism. Rapid unrooﬁng—by tectonic extensional

faulting or erosion or both—leads to steep thermal

gradients near the surface of the Earth and nearly iso-

thermal, retrograde decompression metamorphic paths.

606 Igneous and Metamorphic Petrology