5 Origin of High-Pressure DisorderedMetastable Phases (Lonsdaleite andIncipiently Amorphized Quartz) inMetamorphic Rocks: GeodynamicShock or Crystal-ScaleOverpressure?

Gaston Godard1 2, Rosaria Palmeri3

and David C. Smith4

1Institut de Physique du Globe de Paris, Universite Denis-Diderot, Paris, France2Dipartimento Scienze della Terra, Universita di Siena, Siena, Italy3Museo Nazionale dell’Antartide, Universita di Siena, Siena, Italy4Laboratoire de Mineralogie et Cosmochimie du Museum, Museum Nationald’Histoire Naturelle, Paris, France

5.1 Introduction

Two anomalous metastable phases have been recently reported from ultrahigh-

pressure metamorphic (UHPM) rocks: (i) α-quartz incipiently amorphized under

pressure (IAUP quartz) found in eclogites from Antarctica (Palmeri et al., 2009) and

(ii) lonsdaleite, a hexagonal polytype of sp3-bonded carbon, observed in diamond-

bearing gneiss from the Kokchetav Massif, Kazakhstan (Dubinchuk et al., 2010).

Lonsdaleite has also been reportedly found in a dozen other high-pressure rocks

during the last 30 years (e.g., Golovnya et al., 1977), but these findings, largely

ignored by the scientific community, are doubtful and need to be re-evaluated.

These two phases should not form at equilibrium in metamorphic rocks, as they do

not have a P�T stability field. However, metastable IAUP quartz and lonsdaleite

can form in place of coesite and diamond, respectively, if crystallization of the latter

phases is hampered by kinetics. Besides the UHPM occurrences, IAUP quartz and

lonsdaleite have been observed in impactites and shocked meteorites. If shock is a

necessary process for their creation, then their existence in a supposedly nonshock

environment like UHPM geological units is surprising.

The present chapter is a review on these two metastable phases, which show

many analogies and whose occurrences in UHPM rocks have not received the

Ultrahigh-Pressure Metamorphism. DOI: 10.1016/B978-0-12-385144-4.00004-7

© 2011 Elsevier Inc. All rights reserved.

, Maria Luce Frezzotti

attention that they merit. Its aim is thus to draw attention to these singular phases,

to appraise the reliability of their findings and to enlighten the mechanisms that

could have produced them. We first present and examine the reported findings of

IAUP quartz (Section 5.2) and lonsdaleite (Section 5.3) in UHPM rocks. The

mechanisms that could explain them are subsequently discussed (Section 5.3), and

the conclusion section is devoted to the possible consequences on UHPM research.

5.2 Quartz Incipiently Amorphized Under Pressure

5.2.1 Physical and Crystallographical Background

Pressure-induced amorphization (PIA) in the solid state (i.e., without melting)

involves an initial unstable, low-pressure, crystalline phase (i.e., phase A in

Figure 5.1), whose transformation into a considerably more dense, stable, high-

pressure, crystalline phase (i.e., phase C in Figure 5.1) is kinetically hindered. In

place of such a stable phase (C), the unstable phase (A) transforms into a slightly

more dense, metastable, amorphous state (i.e., phase B in Figure 5.1) that lowers

its free energy (Hemley et al., 1994; see reviews in Sharma & Sikka, 1996; Richet &

Gillet, 1997). PIA of α-quartz was investigated experimentally and theoretically by

many physicists because of its theoretical implications and technological applications

(Hemley et al., 1988, 1994; Kingma et al., 1993a,b; Sharma & Sikka, 1996;

Figure 5.1 Thermodynamic conditions for the formation of metastable phases. Schematic

diagram inspired from the “three-level diagram” of Sharma and Sikka (1996, Fig. 15). GA,

GB, and GC: free energies of phases A, B, and C; italic and uppercase characters refer to

SiO2 and C phases, respectively; dotted vertical lines: diffusion-dependent transitions

impeded by kinetics; full vertical line: diffusionless transition allowed by kinetics.

126 Ultrahigh-Pressure Metamorphism

Richet & Gillet, 1997; Badro et al., 1998). However, PIA has been predicted to be

irrelevant under normal geological pressure�temperature (P�T) conditions, because

the pressures that are needed (e.g.,.12.5 GPa for α-quartz at 300 K) would corre-

spond to extreme depths, and hence to high temperatures at which crystal�crystal

transitions are expected rather than amorphization (Hemley et al., 1994; Sharma &

Sikka, 1996; Richet & Gillet, 1997). Indeed, only the shocks associated with meteor-

ite impacts have so far been found to produce PIA in nature (Goltrant et al., 1992;

Ostroumov et al., 2002).

Densification of quartz under pressure, far inside the coesite and stishovite sta-

bility P�T fields (Figure 5.2), mainly consists of the bending of the Si�O�Si

angle (D’Amour et al., 1979; Levien et al., 1980; Hazen et al., 1989; Chelikowsky

et al., 1990; Glinnemann et al., 1992; Angel et al., 1997; Kim-Zajonz et al., 1999).

At the onset of amorphization, this distortion provokes an irreversible break of

Figure 5.2 P�T phase diagram for SiO2. Modified after Hemley et al. (1994); regular and

italic characters refer to stable and metastable phases, respectively; T.: trydimite; C.:

cristobalite. Solid lines: stable-phase boundaries (A: melting curve of quartz); dotted lines:

metastable extensions of phase boundaries (B: metastable extension of the melting curve of

quartz); dashed area: estimated P�T conditions for the metamorphic peak of the Lanterman

Range rocks (Ghiribelli et al., 2002); dashed line: possible palaeo-geotherm for the

Lanterman Range (Palmeri et al., 2007), assuming a rock density of 3.3; shaded gray area:

progressive amorphization of metastable α-quartz at 300 K (Hemley et al., 1988, 1994;

Hazen et al., 1989; Kingma et al., 1993a; Sharma & Sikka, 1996; Richet & Gillet, 1997);

quartz I/II, coesite I/II: furtive transitions in metastable quartz and coesite at 300 K

(see Hemley et al., 1994).

127UHP Metastable Phases: Shock or Overpressure?

some Si�O links, so both densification and incipient amorphization are two aspects

of the same pressure-driven process. At room temperature, quartz progressively

amorphizes between 12.5 and 40 GPa (e.g., Hemley et al., 1988, 1994; Hazen

et al., 1989; Kingma et al., 1993a,b; Figure 5.2). Using transmission electron

microscopy (TEM), Dell’Angelo (1991) observed an amorphous silica phase along

cracks in quartz-rich aggregates deformed at 600�C, 1025 s21 and a confining P as

low as 1.5 GPa, but this was not confirmed by more recent experiments (Hirth and

Tullis, 1994). PIA can only proceed if Gamorphous silica,Gquartz (i.e., GB,GA in

Figure 5.1), a condition that is fulfilled above 10 GPa by crossing the

metastable extension of the quartz melting curve (B in Figure 5.2) (Hemley et al.,

1988; Badro et al., 1998). Quartz then becomes amorphized in the solid state with

an increasing density, in contrast to the stable part of the melting curve (A in

Figure 5.2) where it melts with a decreasing density. After P relaxation, the lattice

of partly amorphized quartz can retain a little of the densification (i.e., a hysteresis

effect; e.g., Kuznetsov, 2004), together with specific defects that can be identified

by Raman spectrometry (see below).

5.2.2 Evidence from High-Pressure Metamorphic Rocks

Incipient PIA of α-quartz has been found in eclogites from the Lanterman Range

(northern Victoria Land, Antarctica), and identified with in situ Raman spectros-

copy and X-ray microdiffraction using a synchrotron radiation source (Palmeri

et al., 2009). These eclogites formed during the Cambro-Ordovician Ross orogeny

under high-grade conditions (1.6,P, 3.3 GPa and 750, T, 850�C: Palmeri

et al., 2007, and references therein). We summarize here the observations made by

Palmeri et al. (2009).

About a hundred quartz crystals enclosed in garnet and pyroxene were systemat-

ically surveyed. This revealed rare (B5%) monocrystalline quartz inclusions,

50�200 μm in size, exhibiting peculiar Raman spectral features, characterized by

significant discrepancies from typical α-quartz spectra. These are mostly observed

in the inner part of the grains, whereas the outer rim displays the usual α-quartzvibrations (Figure 5.3A). These features can be observed in the same grain without

any change in sample orientation or analytical conditions during spectral data col-

lection. These anomalies consist of the softening and broadening of the 465-cm21

A1 fundamental mode—that may be additionally shifted 1�2 cm21 to lower wave

numbers (i.e., disorder: see Smith et al., 2010)—associated with a marked increase

in intensity of the modes at 267, 402, and 807 cm21 (Figure 5.3A). Furthermore,

one or two additional bands appear at 480�485 cm21, and occasionally at

608 cm21, together with a weak peak at 520�523 cm21. These features are coher-

ent and appear heterogeneously across single grains, as evidenced by Raman map-

ping (Figure 5.4).

The decline in intensity and broadening of the quartz 465 cm21 mode is indicative

of an important structural disordering caused by a greater spread of inter-tetrahedral

128 Ultrahigh-Pressure Metamorphism

Si�O�Si angles (Hazen et al., 1989; McMillan et al., 1992; Ostroumov et al.,

2002). The simultaneous increase in intensity of the 267, 402, and 807 cm21 modes

also supports this interpretation (Resseguier et al., 2003; Figure 5.3A). Meanwhile, the

two bands at about 485 and 608 cm21 correspond to D1 and D2 defects induced by

newly formed isolated ring structures comprising four- and three-membered SiO4

groups, respectively (Sharma et al., 1981; McMillan et al., 1984; Humbert et al.,

1992). None of the crystalline polymorphs of SiO2 contain such defects. They are

reported, however, (i) in quartz partially amorphized at 300 K during static (Hemley

et al., 1988; Kingma et al., 1993b) and shock (Champagnon et al., 1996; Resseguier

et al., 2003) compression (Figure 5.3B), (ii) in silica glass densified at high pressure

Figure 5.3 Raman spectra of partly amorphized α-quartz. (A) Spectra of the inner part(core) and of the outer rim (rim) in a single anomalous quartz inclusion in omphacite

(sample GL1 from Antarctica; Palmeri et al., 2009). Spectroscopic analyses were made at

the same analytical conditions and sample orientation. Star indicates a peak from the resin

mount. (B) By comparison, spectra of one 1-mm-thick quartz crystal unshocked (dotted line)

and shocked at 60 GPa (solid line), from Resseguier et al. (2003).

129UHP Metastable Phases: Shock or Overpressure?

(McMillan et al., 1984; Hemley et al., 1986), and also (iii) in natural quartz limited to

meteoritic impacts (Goltrant et al., 1992; Ostroumov et al., 2002). Finally, the peak at

about 521 cm21 corresponds to the Si�O�Si vibration in four-membered rings of

SiO4 as in coesite (Boyer et al., 1985).

One of the quartz inclusions displaying the above Raman features was investi-

gated further by in situ X-ray microdiffraction using a synchrotron radiation source

(Palmeri et al., 2009). A total of 441 two-dimensional X-ray diffraction patterns

were obtained by scanning the inclusion under a sub-μm diameter monochromatic

X-ray beam, maintaining a constant sample orientation. The patterns display spots

that can be ascribed to a single α-quartz crystal, without any noticeable lattice dis-

orientation. These spots exhibit irregular comet-like shapes (Figure 5.5B and C)

and are about an order of magnitude less intense than those of the host pyroxene,

which show intense circular or elliptic spots (Figure 5.5B) with Gaussian profiles.

As each X-ray exposure was normalized online, the observed intensity differences

indicated that the host pyroxene diffracts X-rays much more intensely than the

quartz inclusion. This suggests that the latter has a relatively low degree of crystal-

linity. Moreover, some of these spots are radially elongated in a direction that

implies a shortening of the corresponding d-spacing by about 2%. This represents a

substantial and irreversible contraction of the lattice with respect to that of normal

α-quartz at 0 GPa (Figure 5.5B and C). Finally, the patterns also display some

weak, hazy “clouds” that can be considered as very weak and broad spots (gray cir-

cles in Figure 5.5A). Their positions do not coincide with planes of α-quartz knownto diffract, but their d-spacing (2.8, 2.7, 2.3, 2.2, 1.9, 1.8, 1.7 A) relate them to coe-

site (2.76, 2.69, 2.33, 2.29, 2.18, 2.03, 1.84, 1.79, 1.71, 1.70 A). Crystal sizes can

Figure 5.4 Raman mapping of a quartz inclusion in omphacite, showing the distribution of

the D1 defect band. Spectral window: 480�490 cm21; same inclusion as for Figure 5.3A;

vertical bar (color in web version) reports relative intensity.

130 Ultrahigh-Pressure Metamorphism

be estimated from the radial breadth of reflections using Scherrer’s equation; the

result suggests that these “clouds,” whatever their origin, are related to lattice fea-

tures whose size is in the order of 10�100 nm (Palmeri et al., 2009).

In summary, in situ Raman spectrometry and X-ray microdiffraction character-

ize the anomalies with respect to α-quartz as (i) a pressure-induced disordering and

incipient amorphization, mainly revealed by new D1 and D2 Raman defect bands

together with (ii) a slight lattice densification, evidenced by d-space shortening,

and (iii) the apparent cryptic development of coesite, 521 cm21 being the wave

number of the main Raman band of coesite. Thus, quartz amorphization (i.e., tran-

sition from A to B in Figure 5.1) was partial, whereas coesite formation (i.e., transi-

tion to C in Figure 5.1) was extremely limited, being apparently restricted to the

formation of cryptic coesite at the nanoscale.

Figure 5.5 X-ray diffraction pattern of the anomalous quartz. (A) Whole pattern. Black

crosses represent the modeled spot positions for α-quartz at 0 GPa and room temperature

(cell parameters from Angel et al., 1997); the angle values are the estimated discrepancies

to the Bragg law (δθ: angle between the real orientation of the diffracting plane and its

orientation if it would strictly obey the Bragg law); gray circles: hazy “clouds” likely

related to coesite (see text); squares: diffraction spots related to the host omphacite.

(B) Detail of spot (202) of α-quartz (estimated discrepancy to the Bragg law: δθ � 0.4�);(C) detail of spot (022) of α-quartz (δθ � 1.2�). Crosses represent the modeled spot

positions for α-quartz between 0 and 6 GPa, at room temperature (cell parameters from

Angel et al., 1997); same crystal as for Figures 5.3 and 5.4; the diffraction pattern was

obtained from a synchrotron radiation source (see text and Palmeri et al., 2009, for the

analytical conditions).

131UHP Metastable Phases: Shock or Overpressure?

Besides impactites, IAUP quartz has only been evidenced, so far, in the above-

mentioned Lanterman Range eclogites, Antarctica (Palmeri et al., 2009), in the

form of monocrystalline inclusions within omphacite. This finding is not doubtful,

since the observation was made in situ and is reproducible. This kind of anomalous

quartz cannot be mistaken for other possible forms of disordered quartz, which

would not show the same Raman spectra and signs of lattice densification. Alpha-

particle irradiation of quartz, for example, has a very weak effect on the Raman

spectra (Krickl et al., 2008), and metamictization of quartz can be achieved only

artificially through high-energy neutron irradiation (Bates et al., 1974).

Similar IAUP quartz probably also exists in similar contexts, where it would

have remained undetected because of its very ordinary appearance and the need for

Raman microspectroscopy. Some micro-inclusions of quartz, for example, included

in zircon from migmatites and showing a weak Raman band of coesite at 521 cm21

(Kobayashi et al., 2008), could be reconsidered in the context of this new approach.

Attention should be paid also to quartz Raman spectra whose main band at

465 cm21 is apparently wide, shifted to higher wave numbers and asymmetric, as

this could result from the superimposition of the 465 cm21 band with the D1 defect

band at 485 cm21. Unfortunately, this D1 defect band can also be hidden by one of

the main garnet bands.

5.3 Lonsdaleite

5.3.1 Physical and Crystallographical Background

Hexagonal sp3-bonded carbon (i.e., the so-called hexagonal diamond) was first synthe-

sized by compressing hexagonal sp2-bonded carbon (i.e., graphite) parallel to the [c]

axis (Bundy & Kasper, 1967; Lonsdale, 1971). This new form of carbon received the

name lonsdaleite when it was found in the Canyon Diablo meteorite (Frondel &

Marvin, 1967). It was soon considered as formed during shock, together with diamond,

either in meteorites (Hanneman et al., 1967; Vdovykin, 1970, 1972; Berkley et al.,

1980; Masaitis et al., 1990; Rubin, 2006; Karczemska et al., 2007) or in impactites

(Masaitis et al., 1972, 1999; Sokhor & Futergendler, 1975; Vishnevskii & Pal’chik,

1975; Hough et al., 1995; Marakushev, 1995; Grieve & Masaitis, 1996; Koberl et al.,

1997; Masaitis, 1998; Gurov et al., 1999; Tsymbal et al., 1999; El Goresy et al., 2003;

Grakhanov, 2005; Kennett et al., 2009; see review in Langenhorst, 2002; Fel’dman

et al., 2007). Lonsdaleite is particularly widespread in the impactites from the Popigai

astroblem (Siberia, Russia: Vishnevskii & Pal’chik, 1975; Koberl et al., 1997;

Tsymbal et al., 1999; El Goresy et al., 2003). The richest lonsdaleite occurrences,

however, are placers where lonsdaleite from presumed impactites have been concen-

trated (Yakutian placers, Russia: Sokhor et al., 1973; Kaminskii et al., 1985; Erlich &

Slonimsky, 1986; Kharlashina & Naletov, 1990; Titkov et al., 2004).

Diamond and lonsdaleite are two polytypes of sp3-bonded carbon (see Merlino,

1997, for a review on polytypism), whose structures have been established by

132 Ultrahigh-Pressure Metamorphism

several authors (Spear et al., 1990; Ownby et al., 1992; Yoshiasa et al., 2003; Wu,

2007; Wang et al., 2008). They are both made of carbon atoms tetrahedrally coor-

dinated, making strong bonds to four neighbors through hybrid sp3 atomic orbitals,

contrary to graphite, in which C atoms are sp2-bonded to three neighbors arranged

in the same plane. Cubic 3C diamond has a face-centered cubic lattice, with

an ABC�ABC�ABC stacking sequence, whereas the hexagonal 2H polytype

(i.e., lonsdaleite) has an AB�AB�AB sequence. The 3C and 2H polytype structures

may combine through a complex intermixture, which gives rise to other polytypes

(e.g., 4H, 6H, 8H and 21R) with larger-period stacking sequences (e.g., ABCB�ABCB�ABCB for 4H). The 2nH polytypes are often considered as being 3C dia-

mond with stacking disorder and/or nanotwinning, which modifies the diamond

properties (hexagonal symmetry, optical anisotropy, greater hardness, shift and

enlargement of the Raman band). The 3C and some of the 2nH polytypes can coexist

in one and the same crystal, but are difficult to distinguish because they share

numerous crystallographic planes with the same d-spacings. Such crystals are fre-

quently labeled “lonsdaleite,” although this name should be, in principle, restricted

to the pure 2H polytype, according to the IMA rules of mineral nomenclature.

Lonsdaleite has been synthesized experimentally at high pressures from graph-

ite, either under shear by compressing graphite crystals parallel to the [c] axis

(Bundy & Kasper, 1967; Lonsdale, 1971; Yagi et al., 1992) or under shock wave

pressure (Trueb, 1968; Babul & Ziencik, 1988; Kurdyumov et al., 1989; Podurets

et al., 1991; Kurdyumov et al., 2000; Xu & Tan, 2003). It is accepted that the

graphite-to-lonsdaleite transition proceeds by martensitic transformation, that is, by

deformation of the graphite crystalline lattice rather than by diffusion and growth

of crystals of the new phase (Kurdyumov et al., 1980; Gorogotskaya et al., 1989;

Britun & Kurdyumov, 2001; Britun, 2002; Britun et al., 2004). Shock loading

applied on diamond (Andreyev et al., 1994; He et al., 2002; Schouwenaars, 2003)

and indentation of a diamond crystal surface (Gogotsi et al., 1997, 1998; Kailer

et al., 1999) can also produce lonsdaleite, through a diffusionless mechanism (e.g.,

Sowa & Koch, 2001). However, lonsdaleite apparently cannot be produced under

isotropic static P, which tends to indicate that, unlike graphite and diamond, lons-

daleite has no P�T field of stability—it is always a metastable phase that appears

only under special conditions (Figure 5.6). Shock or compression pressures favor-

ably orientated could indeed induce the appearance of an anisotropic phase (i.e.,

lonsdaleite) in place of an isotropic one (i.e., diamond), while isotropic pressure

alone cannot favor lonsdaleite, whose theoretical molar volume is virtually identi-

cal to that of diamond (Ownby et al., 1992). Ab initio calculations suggest that dia-

mond is more stable than lonsdaleite under isotropic pressure (Wang & Ye, 2003),

but would become unstable and transform into lonsdaleite under deviatoric com-

pressive stress with a critical mean stress σm of 26 GPa (Wen et al., 2008;

σm5 [σ11σ21σ3]/3); the value of this parameter is unknown for the graphite-

to-lonsdaleite transition. Although lonsdaleite has no known P�T stability field, it

is established that it transforms instantly into cubic diamond at very high tempera-

tures, beyond zone A of Figure 5.6 (Bundy et al., 1996). Below this zone, at lower

133UHP Metastable Phases: Shock or Overpressure?

T and P, the transformation becomes sluggish and lonsdaleite may remain in the

metastable state under conditions of relatively low temperatures.

The above experimental and theoretical data indicate that lonsdaleite is a

metastable phase whose lattice structure and molar volume are very close to those

of 3C diamond, so it is obviously much less unstable than graphite at ultrahigh

pressures (i.e., Gdiamond,Glonsdaleite{Ggraphite; cf GC,GB,GA in Figure 5.1).

Moreover, the graphite-to-lonsdaleite transition is a diffusionless martensitic trans-

formation. Therefore, metastable lonsdaleite can form from graphite at ultrahigh

pressures instead of 3C diamond, if the formation of the latter is impeded by kinet-

ics (see Figure 5.1). This situation is similar to that of IAUP quartz, which probably

explains why these two phases occur in the same contexts (i.e., impactites, shock

experiments, UHPM rocks). However, there are apparently two differences: (i)

lonsdaleite needs a deviatoric stress to form from graphite (it is not known that

amorphization of quartz needs a deviatoric stress) and (ii) it can also form from 3C

diamond under strong deviatoric stresses (.26 GPa, according to Wen et al.,

2008).

Figure 5.6 P�T phase diagram for carbon. Modified after Bundy et al. (1996); regular and

italic characters refer to stable and metastable phases, respectively. Solid lines: stable-phase

boundaries; dashed line: possible geotherm for a subduction zone, assuming a rock density

of 3.3; the A dotted zone marks the threshold of fast experiments that convert metastable

phases (graphite, lonsdaleite), however generated, into cubic diamond—it can be considered

as the uppermost limit for the metastability of lonsdaleite and graphite; B: transformation of

graphite into retrievable hexagonal sp3-bonded carbon (i.e., lonsdaleite?); C: upper end of

shock-quench cycles that convert graphite to lonsdaleite; D: upper end of shock-quench

cycles that convert graphite to cubic diamond.

134 Ultrahigh-Pressure Metamorphism

5.3.2 Evidence from High-Pressure Metamorphic Rocks

Lonsdaleite was first described in 1977 by Golovnya et al. (1977) in eclogites from

the Kola Peninsula (Sal’nyye Tundry region) and glaucophane-bearing eclogites

from Southern Urals (Shubino; Maksyutov complex). It reportedly occurred as

yellowish semitransparent grains (0.1�0.5 mm) with a pseudohexagonal habit.

X-ray diffraction data were included and show d-spacings either typical of lonsda-

leite (2.17, 1.93�1.95 A) or shared by diamond and lonsdaleite (2.07, 1.26,

1.08 A). Unfortunately, the finding was poorly documented, without any precise

sample location; moreover, lonsdaleite was not observed in situ but separated after

rock grinding, so contamination cannot be absolutely ruled out. Kuzovkov (2001)

also mentioned the presence of lonsdaleite, together with diamond and coesite,

once more in glaucophane-bearing eclogites from the Maksyutov complex (Urals),

and proposed a highly disputable geodynamic model considering the Urals belt as a

gigantic impact structure.

Shumilova et al. (2002) reported crypto-crystalline lonsdaleite, together with

diamond and other unusual carbon phases such as chaoite, in graphite flakes of

Laplandian granulites from the Kola Peninsula; these various carbon phases were

identified with their d-spacings obtained by X-ray diffraction. The same authors

also mentioned several poorly defined phases of sp-, sp2- and sp3-bonded carbon in

schists underlying the Khabarnin ophiolite massif (Southern Urals) and in quart-

zites and schists of the Nerkayus metamorphic complex (Subpolar Urals).

Recently, Dubinchuk et al. (2010) described lonsdaleite in garnet�biotite gneiss

of the Kumdykol diamond deposit (Kokchetav Massif, Kazakhstan). Evidence for

lonsdaleite is based on electron diffraction patterns that show reflections whose

d-spacings are effectively distinctive of lonsdaleite (1.93 and 1.50 A). The latter is

reported to form a 0.3-μm-thick corona around a graphite flake.

In all the above occurrences, lonsdaleite was reported as formed from graphite.

The evidence is particularly strong in the latter case, where lonsdaleite apparently

occurs as a corona around graphite.

Vinokurov et al. (1998) also reported “the probable formation of hexagonal

lonsdaleite layers in the crystal structure of octahedral diamonds” from peridotite

of a kimberlite diatreme of the Liaoning province (northeast China). Their assump-

tion was supported by X-ray diffraction bidimensional patterns that show a pseudo-

hexagonal distribution of the reflections; surprisingly, they did not report d-spacing

values. In this case, the “lonsdaleite layers” would have formed in a pre-existing

3C diamond crystal, the octahedral shape of which was preserved.

Finally, lonsdaleite is suspected to occur in μm-sized diamond inclusions of a

few other UHPM rocks, either within garnet (Rhodope, Greece: Perraki et al.,

2006) or within zircon (Straumen, Norway: Godard et al., 2004, 2005; Kokchetav

Massif, Kazakhstan: Smith et al., 2010). Portions or the totality of these micro-

inclusions show Raman spectra with the band at 1332 cm21—known to correspond

to sp3 bonding of carbon—significantly weaker, wider, and shifted to smaller wave

numbers, in comparison with normal diamond. These Raman features are typical of

lonsdaleite and other hexagonal polytypes, whose band is shifted to 13206 5 cm21

135UHP Metastable Phases: Shock or Overpressure?

(Figure 5.7; Nemanich et al., 1988; Knight & White, 1989; Bhargava et al., 1995; Wu

& Xu, 1998; Karczemska et al., 2007; see review in Smith & Godard, 2009; Smith

et al., 2010). However, some of these referenced spectra, with widely scattered wave

numbers of Raman bands, may not derive from lonsdaleite but from disordered dia-

mond with a high density of lattice defects (Smith et al., 2010). Thus, doubt still exists

over whether diamond is disordered or intermixed with lonsdaleite—options that

moreover could all result from the same process of lattice deformation.

Lonsdaleite is difficult to distinguish from 3C diamond, to which it can be asso-

ciated at the nanoscale. Care must be exercised also to exclude any other possible

disordered form of sp3-bonded carbon. In particular, α-particle irradiation of dia-

mond in contact with a radioactive mineral such as zircon can generate a few

Figure 5.7 Raman spectrum of diamond compared to those of “lonsdaleite.” Raman spectra

of standard diamond (top) and of hexagonal sp3-bonded carbon from Popigai (“lonsdaleite”

samples from the Museum National d’Histoire Naturelle, Paris; from bottom to top: Pop-10,

Pop-60, and Pop-90ptb; see Smith et al., 2010, for the characterization of sample Pop-10 by

X-ray diffraction). Analytical conditions: 325 nm, RENISHAW INVIA, 100�10,000 cm21,

no baseline correction, smoothed 15 points. Note the four very wide, very weak “massifs” at

6600, 7300, 8100, and 8900 cm21 (and also three even weaker ones to their left) that occur

in all three lonsdaleite spectra but also in the standard diamond, such that all are

luminescence phenomena from the INVIA. The diamond spectrum reveals peaks at 1332

(very intense), 2472, and 6716 cm21. The inset zooms on the sp3 bands to display the

diamond centered at 1332 cm21 and the lonsdaleites at 1322 cm21 for the bottom two.

136 Ultrahigh-Pressure Metamorphism

lattice defects (Campbell & Mainwood, 2000) that have some effects on the Raman

band. This yields a resulting spectrum comparable to that of lonsdaleite (Orwa

et al., 2000) in showing the triple tendency of decreased intensity, decreased wave

number, and increased bandwidth of the Raman band (Figure 5.7; e.g., Nasdala

et al., 2008; Smith et al., 2010). Furthermore, nm-sized diamond particles produce

similar effects (Yoshikawa et al., 1993). Polishing of diamond may also damage

the crystal surface, and so have a slight effect on Raman spectra (Knight & White,

1989; Hird et al., 2007). However, irradiation also generates four new Raman peaks

(Orwa et al., 2000) unknown in lonsdaleite spectra, which can help to distinguish

between lonsdaleite and metamictized (i.e., irradiated) diamond. Indeed, Smith

et al. (2010) argue that diamond at Kokchetav (Kazakhstan) has been metamictized

because of the presence of precisely these four Raman peaks.

In summary, lonsdaleite has been reported in several UHPM rocks, but the find-

ings are poorly documented and solid proof is lacking in some cases. Graphite

was, in several cases, totally transformed into lonsdaleite (i.e., the transition from A

to B in Figure 5.1); the existence of intermediate forms between lonsdaleite and dia-

mond indicates a partial transition to diamond (i.e., transition from B to C in

Figure 5.1). When considering the list of natural occurrences of lonsdaleite, it would

be prudent to bear in mind that lonsdaleite, unlike other mineral species, is not easy

to identify definitively, because Raman spectra can be ambiguous and a transition

toward cubic 3C diamond is common, Investigation by in situ X-ray microdiffrac-

tion and/or TEM can provide definitive data by highlighting of specific d-spacings

of lonsdaleite or other hexagonal polytypes. Unfortunately, carbon micro-inclusions

are generally too tiny to be easily investigated by X-ray diffraction and TEM.

5.4 Discussion

Quartz incipiently amorphized under pressure (IAUP quartz) and lonsdaleite are

both disordered and metastable phases formed under ultrahigh pressures and/or

deviatoric stresses. They have been almost totally ignored so far, probably because

they are difficult to recognize and can also be mistaken for other disordered forms

of silica and carbon.

PIA of quartz and the formation of lonsdaleite from graphite can only proceed if

the two following conditions are fulfilled (Figure 5.1):

i. Quartz and graphite should be brought far outside their P�T stability field, up to ultra-

high pressures at which they can be transformed, at least partially, into amorphous silica

and lonsdaleite, respectively (i.e., GA.GB in Figure 5.1). The condition

Gquartz.Gamorphous silica is fulfilled by crossing the metastable extension of the quartz

melting curve (B in Figure 5.2; see Section 5.2.1), but this extension is not known with

great accuracy. At 800�C, it would take place at P.B10 GPa, since experiments have

shown that quartz PIA initiates between 12.5 and 15 GPa at room temperature (see

Figure 5.2 and Section 5.2.1). As for carbon, the P�T conditions for Ggraphite.Glonsdaleite

137UHP Metastable Phases: Shock or Overpressure?

are unknown, but they obviously correspond to a portion of the diamond P�T stability

field (Figure 5.6), since graphite is unstable in these conditions. Moreover, deviatoric

stress seems also necessary to form lonsdaleite from graphite and also from diamond (see

Section 5.3.1). In all cases, the pressures necessary to produce IAUP quartz and lonsda-

leite are far above any reasonable geotherm (Figures 5.2 and 5.6), so a nonlithostatic

component to the pressure is required. Most of the other UHPM minerals (i.e., garnet,

pyroxene, coesite, and rutile) are not known to amorphize or alter under pressure, most

likely because they are stable up to extremely high pressures (i.e., the condition GA.GB

of Figure 5.1 is never fulfilled). White micas and sulfides, however, could also be

affected by PIA (Sharma & Sikka, 1996).

ii. Unfavorable kinetics should prevent the formation of stable phases like coesite and dia-

mond (stable phases C of Figure 5.1). This can occur in two cases: (a) during a long

high-pressure event, at low T, when diffusion is hence too slow for enabling any

solid�solid transitions (i.e., T,B500�C for quartz; Perrillat et al., 2003); (b) at medium

or high T, during a very short high-pressure event (i.e., a shock); in the case of the

Lanterman Range eclogite, this event should normally have been of less than 3 h at the

peak T (B800�850�C), otherwise quartz should have been totally transformed into coe-

site (see discussion in Palmeri et al., 2009, and the kinetic law for the quartz-to-coesite

transition in Perrillat et al., 2003); lonsdaleite, however, can easily survive during short

events up to T � 2000 K (i.e., up to the A dotted zone of Figure 5.6). These two situa-

tions (a and b) can combine in the case of a shock at low temperature.

These general considerations lead us to consider two main possible mechanisms

for the formation of IAUP quartz and lonsdaleite in UHPM rocks, namely shock

and overpressure. While considering these mechanisms, keep in mind that both

generally occur under deviatoric stress. Under such conditions, the main stress (σ1)

can control the formation of high-pressure phases, either stable or metastable,

instead of the confining P (σ3) or mean stress (σm5 [σ11σ21σ3]/3). This is

strongly suggested by the experiments on deformation of quartz aggregates by

Hirth and Tullis (1994), which produced coesite along quartz grain boundaries ori-

ented perpendicular to σ1, at 500�700�C, while the mean stress (σm, 1.7 GPa)

and confining P (σ3, 1.3 GPa) were below the quartz�coesite transition, whereas

the maximum stress was above the transition (σ1. 2.8 GPa).

5.4.1 Shock

Lonsdaleite and IAUP quartz are well known in impactites (see Sections 5.2.1 and

5.3.1), where they obviously formed during a shock. One can wonder what could

generate such shocks in UHPM rocks.

Earthquakes (Smith, 1988; Austrheim & Boundy, 1994; Austrheim et al.,

1996; John & Schenk, 2006) and tectonic pulses (Camacho et al., 2005; Kelley,

2005) have been occasionally discussed in the context of eclogite formation.

However, the stresses related to an earthquake seem insufficient to explain the

development of such minerals. Rocks are approximately elastic at the seismic-wave

frequencies, and the elastic parameters involved during an earthquake (Young’s

modulus, differential stress, etc.) are from 10 to 100 MPa as orders of magnitude

(Ranalli, 1995; Kanamori, 2004; Stacey & Davis, 2008). Stress can locally

138 Ultrahigh-Pressure Metamorphism

concentrate at the crystal scale, because of crystal anisotropy or obstacles to stress-

wave propagation for example, but it is hardly believable that it could increase of

several orders of magnitude, up to B10 GPa. Paradoxically, if earthquakes were

able to produce high-pressure phases, these would be very common in the whole

lithosphere.

Other causes with strong inelastic effects can be envisaged. The well-known

kimberlite explosive eruptions, or the hypothetical “Verneshot” explosions (Morgan

et al., 2004), could be invoked, for example, in the case of the lonsdaleite report-

edly found in the Liaoning diatreme (Vinokurov et al., 1998). The hypothesis of a

meteorite impact can also be considered. Although the rocks under consideration do

not show any traces of cataclasis and are clearly not impactites, the shock wave

could have produced some effect at significant distances from the point of impact,

even as far as the antipodes according to the antipodal model (Hagstrum, 2005).

To explain the presence of lonsdaleite in metamorphic rocks, Vishnevskii and

Raitala (2000) envisaged that it was inherited from Precambrian impactites and

reworked in sediments, where it would have survived the subsequent metamor-

phism. Indeed, impact diamond or lonsdaleite crystals with preserved disorder fea-

tures have been observed in rocks of the Sudbury impact structure metamorphosed

in greenschist-facies conditions (Ontario, Canada; Masaitis et al., 1999), and it

seems that lonsdaleite can survive as a metastable phase at rather high temperatures

(see Figure 5.6). IAUP quartz, however, would easily recover at high or medium

temperatures, so it could hardly survive a high-grade metamorphism (P. Richet,

personal communication). Moreover, such a reworking could be imagined for meta-

sediments, but it is much unlikely for eclogite and other meta-igneous rocks.

5.4.2 Overpressure

Pressure with a strong nonlithostatic component, generally called “overpressure,”

could cause the formation of metastable IUAP quartz or lonsdaleite, provided that

this overpressure is sufficiently intense and operates during a relatively brief span

of time and/or at a rather low temperature to prevent the crystallization of coesite

or cubic diamond, respectively.

Overpressure extending on a regional scale has been invoked to explain the for-

mation of UHPM rocks in mountain belts (see review in Smith, 1988), and a few

recent models suggest the local existence in subduction zones of a nonlithostatic

P component that can reach 0.3�3.0 GPa, depending on the model (Mancktelow,

1993, 1995, 2008; Petrini & Podladchikov, 2000; Raimbourg & Kimura, 2008;

Vrijmoed et al., 2009; Li et al., 2010). However, the nonlithostatic component is

limited relative to the whole pressure (,20% according to Li et al., 2010), otherwise

the lithosphere strength would be unable to maintain it. Thus, the inferred overpres-

sure seems insufficient to explain the formation of the above metastable phases, par-

ticularly IAUP quartz, which needs a total pressure of more than 10 GPa (see

above).

Overpressure at the crystal scale seems more adequate. It could result from a

volume mismatch during the retrograde P�T evolution between a micro-inclusion

139UHP Metastable Phases: Shock or Overpressure?

(quartz, graphite, diamond) and its rigid and resistant container (zircon, garnet,

pyroxene, and diamond) (Gillet et al., 1984; van der Molen & van Roermund,

1986; Guiraud & Powell, 2006). Raman spectroscopy and X-ray diffraction have

demonstrated that some quartz and/or coesite micro-inclusions enclosed within

unfractured garnet, zircon, or diamond have preserved a present-day overpressure

as high as 1.6 GPa (Korsakov et al., 2007, 2009), 1.9�2.3 GPa (Parkinson &

Katayama, 1999; Parkinson, 2000), 2.4 GPa (Ye et al., 2001), and 3.6 GPa

(Sobolev et al., 2000). Added to a substantial lithostatic component, such pressures

could contribute to reach the ultrahigh pressures required to generate IAUP quartz

and lonsdaleite. This hypothesis is also supported by the microstructures of these

phases, which apparently appear as micro-inclusions when observed in situ. The

anisotropy of both enclosing and enclosed minerals could play an important role,

since some reciprocal lattice orientations between inclusion and container can

induce a significant volume mismatch and others not. Anisotropy could also induce

the deviatoric stress apparently necessary for the lonsdaleite formation.

5.5 Conclusion

Incipiently amorphized α-quartz and lonsdaleite are anomalous phases that cannot

crystallize at equilibrium in metamorphic rocks. They form in place of coesite/stisho-

vite and diamond, respectively, if crystallization of the latter minerals is impeded by

kinetics; for example, during a shock or at low temperatures. Besides impactites,

these metastable phases have been observed thus far in a few UHPM rocks, but they

could be more common. Indeed, they remain difficult to put in evidence, because (i)

they can be identified only by high-resolution in situ techniques (e.g., Raman spec-

trometry, X-ray microdiffraction, and/or TEM), (ii) they are preserved in narrow

μm-sized zones, and (iii) they show a spatial transition toward normal α-quartz and

diamond. For these reasons, they have almost completely escaped notice so far.

The purpose of this chapter is thus to draw the attention of the scientific com-

munity to these phases, rather than to provide an unambiguous explanation on their

origin, which is still premature (see above). Only new observations could answer

the following questions:

i. Whether these phases are really occasional, linked to a local and exceptional phenome-

non, or exist more commonly in UHPM rocks, reflecting a fairly general metamorphic

mechanism.

ii. Whether they are restricted to μm-sized inclusions within resistant containers (zircon, gar-

net, omphacite, or diamond), which could mean that their development and preservation

are favored by intracrystalline overpressures at the inclusion scale, or they are also inter-

granular, which would tend to favor the hypothesis of their formation during a shock.

iii. Whether they only occur in UHPM rocks or also exist in other metamorphic rocks.

iv. Whether other similar metastable phases occur in UHPM rocks.

Whatever the answers to the above questions, the occurrence of such high-

pressure disordered metastable phases strongly supports the hypothetical role of

140 Ultrahigh-Pressure Metamorphism

nonlithostatic P in UHPM rocks, at least either during short events (i.e., shocks) or

on a minute scale (i.e., inclusion-scale overpressure).

Acknowledgments

We are grateful to R.J. Davies for his help at the synchrotron, to S. Wallis for the editorial

work, and to G. Ranalli and P. Richet for stimulating discussions on physics of earthquakes

and thermodynamics of pressure-induced amorphization, respectively.

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