stability, metastability, and elastic properties of a...
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JOURNAL OF GEOPHYSICAL RESEARCH: SOLID EARTH, VOL. 118, 1–13, doi:10.1002/jgrb.50360, 2013
Stability, metastability, and elastic properties of a dense silicapolymorph, seifertiteB. Grocholski,1 S.-H. Shim,2 and V. B. Prakapenka3
Received 23 July 2012; revised 20 June 2013; accepted 29 August 2013.
[1] Dense silica polymorphs with sixfold coordinated Si have been found in SNC andlunar meteorites and may be important minerals for silica-rich components in the lowermantle. However, the stable crystal structure in the lower mantle and properties of densesilica remain controversial. Under stable heating and quasi-hydrostatic stress conditions,we found that the CaCl2 type undergoes a phase transition to the ˛-PbO2 type (seifertite)at 130–140 GPa and 2500 K. Our data suggest that this phase transition occurs at agreater depth than the perovskite! postperovskite transition in the lowermost mantle.The molar volume measured at 1 bar is the smallest among the reported silicapolymorphs, therefore having the highest calculated density and in excellent agreementwith recent first-principles calculations. The greater molar volume of seifertite found inthe shergottite meteorite and previous experiments supports a metastable synthesis of thephase outside its stability field. Our data combined with the Hugoniots of silicapolymorphs also rule out the possibility of the formation of seifertite in the meteoritewithin its stability field. We found very little change in bulk sound speed across theCaCl2-type! seifertite transition. If shear wave velocity decreases at the transition toseifertite as suggested by some computational studies, this silica transition may providean alternative explanation for the discontinuities with a shear wave velocity decreasefound at depths greater than the D00 discontinuity.Citation: Grocholski, B., S.-H. Shim, and V. B. Prakapenka (2013), Stability, metastability, and elastic properties of a dense silicapolymorph, seifertite, J. Geophys. Res. Solid Earth, 118, doi:10.1002/jgrb.50360.
1. Introduction[2] Silica holds a special place in geophysics and plane-
tary sciences as it is one of the most abundant componentsin the Earth and other terrestrial planets. The equilibriumphase diagram of silica is becoming clear from high pres-sure experiments, computations, and studies on AX2 analogcompounds [e.g., Haines et al., 1996; Haines and Leger,1997; Prakapenka et al., 2003a]. The transition from four-fold coordinated (coesite) to sixfold coordinated (stishovite,rutile-type structure) Si-O occurs at �7 GPa [Zhang et al.,1996]. The stishovite structure distorts slightly to the CaCl2-type structure at �50 GPa [Kingma et al., 1995; Andraultet al., 1998a; Ono et al., 2002]. A few studies have sug-gested that the CaCl2 type may undergo a phase transitionto a mineral named seifertite with the scrutinyite (˛-PbO2)structure within mantle pressures [Dubrovinsky et al., 2001;
1Department of Mineral Sciences, National Museum of Natural History,Smithsonian Institution, Washington, DC, USA.
2School of Earth and Space Exploration, Arizona State University,Tempe, Arizona, USA.
3GeoSoilEnviroCARS, University of Chicago, Argonne, Illinois, USA.
Corresponding author: B. Grocholski, Department of Mineral Sciences,National Museum of Natural History, Smithsonian Institution, 10th &Constitution Ave., Washington DC 20560, USA. ([email protected])
©2013. American Geophysical Union. All Rights Reserved.2169-9313/13/10.1002/jgrb.50360
Murakami et al., 2003; Tsuchiya et al., 2004; Oganov et al.,2005; Driver et al., 2010] and persist until the pyrite struc-ture (6 + 2 Si-O coordination) is stable above �270 GPa[Kuwayama et al., 2005].
[3] Seifertite is an important mineral for both under-standing the lower mantle of the Earth [Hirose et al.,2005; Grocholski et al., 2012] and is perplexing for itsappearance far outside the stability field in SNC-type andlunar meteorites [Stöffler et al., 1986; Sharp et al., 1999;Boctor et al., 2003; Aoudjehane et al., 2005; El Goresy et al.,2008; Miyahara et al., 2013]. Silica may not be a domi-nant mineral in the lower mantle as homogenized mantlecompositions, such as pyrolite, may be silica undersatu-rated [Ringwood, 1982]. However, Bina [2010] suggestedthe possible existence of free silica even in the pyroliticmantle as polycrystalline-armored relics with seismicallydetectable scale. Several recent studies have suggested thepossibility of a heterogeneous lower mantle including sil-ica oversaturated components that would provide diversityto the mineralogy of the region [Nakagawa et al., 2010;Grocholski et al., 2012].
[4] The reported transition pressure of silica to seifertitefrom the CaCl2-type structure varies over a wide range andthe existence of the transition in the mantle has remainedunclear. Dubrovinsky et al. [1997] reported a phase transi-tion to a seifertite-like phase at �70 GPa, while Murakamiet al. [2003] reported much higher pressure for the
1
GROCHOLSKI ET AL.: PROPERTIES OF SEIFERTITE
Table 1. Experimental Conditionsa
Culet Pressure Temperature Total HeatingPressure Pressure Size Data Range Range Duration Phase
Samples Standard Medium (�m) Used (GPa) (K) (min) Observedc
Pure SilicaAr-1 Au Ar 75 �0, eos 139–146 1900–2600 90 sftAr-3 Au Ar 100 126 3100 20 sft + CTNe-5 Ne Ne 100 152 2700–2900 120 sftNe-7 Au Ne 150 �0, eos 102 2020 21 CT
Al-Bearing Silicab
Ar-2Al Au Ar 75 �0, eos 125–158 1900–2900 130 sft + trace CTNe-6Al Au Ne 150 �0, eos 120 2500–2700 120 CTAr-4Al Pt Ar 200 61–75 1700–2500 210 CT
a�0: density calculated from volume measurements at 1 bar (Table 2), eos: P-V measurements during decompression (Figure 5).b�10 mol % Al2O3.csft: seifertite, CT: CaCl2 type.
CaCl2-type! seifertite phase transition at 121 GPa. Theseearlier studies were performed without pressure transmittingmedium and thermal insulation in the diamond-anvil cell,which can result in severe deviatoric stresses and thermalgradients. The free-energy differences between competingsilica phases can be small [Teter et al., 1998], and hence,the formation of metastable phases under nonequilibriumconditions is particularly problematic. A recent study withquasi-hydrostatic conditions and thermal insulation [Shiehet al., 2005] failed to detect the phase transition to seifertiteup to �130 GPa using the same pressure scale as Murakamiet al. [2003]. However, computational studies suggest thatthe phase transition should occur at 90–120 GPa [Karkiet al., 1997a; Tsuchiya et al., 2004; Oganov et al., 2005;Driver et al., 2010].
[5] Seifertite has been found on the surface of theEarth in meteorites of Martian and lunar origin (Shergotty,Zagami, NWA 4734), both of which have high (30–80 GPa)estimated peak shock pressures [Stöffler et al., 1986; Sharpet al., 1999; Boctor et al., 2003; Aoudjehane et al., 2005; ElGoresy et al., 2008; Miyahara et al., 2013]. However, theuncertainties in the stable P-T conditions for seifertite makeit difficult to understand how this mineral is synthesized inthese meteorites. In addition, the calculated density from themeasured volume of seifertite from the Shergotty and NWA4734 meteorites [Dera et al., 2002; Miyahara et al., 2013]and from cold decompression of cristobalite [Tsuchida andYagi, 1990; Dubrovinsky et al., 2001] is as much as 2%lower than the recent computational calculations [Oganovet al., 2005; Driver et al., 2010]. If the density difference isconverted into pressure based on equation of state of densesilica, it would be equivalent to 9 GPa.
[6] We have measured the seifertite phase transition insilica at pressures up to 152 GPa and temperatures of2000–3500 K in the laser-heated diamond-anvil cell usinginsulating compressed noble gas media. We also were ableto decompress seifertite synthesized at stable P-T condi-tions (�140 GPa at �2500 K) to room pressure, allowingus to constrain the equation of state of seifertite with a 1 barvolume (V0) and calculate a 1 bar density (�0).
2. Experimental Method[7] A total of seven crystalline samples were synthesized
at high pressure and temperature. Samples of pure (99.8%)
amorphous silica or 10 mol % Al2O3-bearing amorphous sil-ica were mixed with 15 wt % Au, 10 wt % Pt, or 30 wt %Fe, compressed into platelets, and surrounded by condensedinert noble gas (Ar or Ne) for thermal insulation and quasi-hydrostatic stress environment (experimental setup can befound in Table 1). Argon was cryogenically loaded at theMassachusetts Institute of Technology and neon was gas-pressure loaded at the Advance Photon Source [Rivers et al.,2008]. The Al-bearing starting materials were synthesizedby aero-levitation [Tangeman et al., 2001].
[8] Diamond cells equipped with anvils having 75, 100,150, or 200 �m culets were used with initial samplechambers of 35, 50, 90, or 120 �m in diameter, respec-tively. We conducted X-ray diffraction in situ at high P-Tin the double-side laser-heated diamond-anvil cell at theGSECARS beamline 13IDD at the Advanced Photon Source[Prakapenka et al., 2008]. All samples were heated to T >2000 K for 20–210 min after reaching target pressure withX-ray diffraction collected using a marCCD detector every�2–3 min to track crystal growth during heating (Table 1).
[9] Diffraction patterns were collected during decompres-sion at 3–5 GPa steps until catastrophic decompression ofthe sample occurred for the equation of state at room tem-perature (Table 2). Since the beam size was only 3–5 �m,the sample was oscillated ˙1–3 �m during collection ofdiffraction patterns to improve statistics and generate moreuniformly smooth diffraction rings. This technique is usedfor the measurements at 300 K, but not for the measurementsduring laser heating. Diffraction patterns were integratedusing the Fit2d image processing program [Hammersleyet al., 1996] with lattice parameters and error calculatedusing the UnitCell program [Holland and Redfern, 1997](Tables 3 and 4). Peak positions and background sub-traction from the 1-D integrated patterns were determinedwith the home-written IDL fitting programs, XPEAKPOand XPEAKFIT (www.public.asu.edu/~sshim5/Dan_Shim/Softwares.html). A total of four samples were recoveredat room conditions, allowing for measurement of the roompressure volume (V0). Densities were then calculated fromthe volumes assuming pure silica or 10 mol % Al2O3-bearing SiO2. Room pressure sample recovery from Mbarpressure was accomplished by precompressing the rheniumgaskets to 40 GPa, well-centered circular sample chambersdrilled with an EDM, and slow release of pressure using
2
GROCHOLSKI ET AL.: PROPERTIES OF SEIFERTITE
Table 2. Equation of State-Fitting Results With Room Pressure Density (�0) Calculated From V0Measured at 1 Bar and 300 K
V0(Å3) K0 (GPa) K00 �0(g/cm3)
Seifertite, This WorkSiO2, all data 91.66(6) 322(2) 4a 4.355(3)SiO2, P > 70 GPa 91.66(6) 337(8) 3.5(2)Al-SiOb
2 92.20(8) 322(2) 4a 4.329(4)
Seifertite, Shergotty, and NWA 4734Dera et al. [2002] 92.92 4.295Miyahara et al. [2013] 91.96 4.340
Seifertite, ExperimentDubrovinsky et al. [2001] 93.52(14) 313(5) 3.43(11) 4.267(6)
Seifertite, TheoryOganov et al. [2005] 91.12 324.4 4.23 4.380Driver et al. [2010] 91.76 329 4.0 4.350
CaCl2 Type, This WorkSiO2, all data 46.63(3) 317(3) 4a 4.279(3)Al-SiOb
2 47.249(3) 298(2) 4a 4.223(3)
CaCl2 Type, ExperimentAndrault et al. [2003] 46.63(2) 334(7) 4a 4.279(2)
CaCl2 Type, TheoryOganov et al. [2005] 47.15 258 4.6 4.232Driver et al. [2010] 47.13 305 3.7 4.235
Stishovite, ExperimentAndrault et al. [2003] 46.51(1) 310(1) 4.6(2) 4.290(1)
Disordered Dense Silica, ExperimentDubrovinsky et al. [2004] 4.286(3)
Postquartz, ExperimentHaines et al. [2001] 4.29(4)
Baddeleyite, ShergottyEl Goresy et al. [2000] 4.30(2)
aFixed during fitting.b� 10 mol % Al2O3.
small turns of the set screws located on the piston side of thesymmetric diamond cells.
[10] The metals mixed in our glass starting material wereused for laser coupling, with gold and platinum also act-ing as internal pressure standards [Tsuchiya, 2003; Holmeset al., 1989]. The laser spot size is roughly 20 �m and theerror on the temperature is 150 K, about twice the variationwe obtain from the gray body fits over the heating dura-tion. Samples did not change in optical appearance after eachheating cycle or upon decompression, indicating no drasticchange in visible light absorption characteristics between thedensified glass, seifertite, and the CaCl2-type structure. Weuse neon as a pressure standard [Dewaele et al., 2008] inour sample mixed with iron as a laser coupler (Ne-5) sincethere is the possibility of carbon contamination of the metal[Prakapenka et al., 2003b] (Table 1).
3. Results3.1. Stable Pressure-Temperature Conditions
[11] A pure SiO2 glass (Ar-1) was heated at 143 GPa and2137 K. New diffraction peaks appeared after �10 min ofheating and they persisted in subsequent heating (Table 1and Figures 1 and 2c). Through first-principles method,
Teter et al. [1998] showed that four different polymorphswith sixfold coordinated Si become energetically favor-able over the CaCl2 type at P �100 GPa. We calculatedthe diffraction patterns of the four polymorphs based onthe structural data in Teter et al. [1998] (Figure 3). Thenew diffraction peaks are well explained by the orthorhom-bic ˛-PbO2-type (seifertite) structure (Figure 3). Absenceof any major diffraction lines between 0.35 and 0.4 Å–1
(between 2.85 and 2.5 Å) in the measured patterns, otherthan nitrogen (see below), rules out the possibility of theother polymorphs, i.e., postquartz, NaTiF4 type, and SnO2type (Figure 3). The weaker diffraction lines at higher anglesare also explained well by the seifertite structure.
[12] The other peaks in the diffraction patterns can beassigned to those from the rhenium gasket, gold internalpressure standard, argon pressure medium, and the cubic-gauche phase of nitrogen (cg-N) [Eremets et al., 2004](Figure 3a). Tails from the X-ray beam combined withthe high scattering cross section of the Re gasket lead toRe peaks in the diffraction patterns for sample chambers<50 �m. The cg-N phase is a rare single-bonded struc-ture formed from small amounts of nitrogen captured duringcryogenetic loading of Ar (this will be described else-where). The diffraction peaks which can be assigned to thecg-N phase were not observed in the Ne-loaded samples
3
GROCHOLSKI ET AL.: PROPERTIES OF SEIFERTITE
Tabl
e3.
Latti
cePa
ram
eter
s(a,
b,c)
and
Uni
tCel
lVol
umes
(V)f
orA
llD
iffra
ctio
nPa
ttern
soft
heSe
iferti
tePh
asea
Pure
Silic
aSi
lica
With
Alu
min
a
P(G
Pa)
a p(Å
)a(
Å)
b(Å
)c(
Å)
V(m
ol/c
m3 )
P(G
Pa)
a p(Å
)a(
Å)
b(Å
)c(
Å)
V(m
ol/c
m3 )
04.
073˙
0.00
35.
026˙
0.00
34.
477˙
0.00
313
.80˙
0.01
04.
082˙
0.00
35.
037˙
0.00
44.
484˙
0.00
313
.88˙
0.01
38.6˙
0.8
3.88
2˙
0.00
33.
945˙
0.00
44.
880˙
0.00
44.
359˙
0.00
312
.63˙
0.01
65.1˙
1.6
3.80
4˙
0.00
23.
850˙
0.00
44.
825˙
0.00
54.
300˙
0.00
512
.02˙
0.02
46.7˙
1.1
3.85
5˙
0.00
23.
894˙
0.00
84.
862˙
0.01
04.
345˙
0.00
612
.38˙
0.02
67.1˙
1.6
3.79
8˙
0.00
23.
849˙
0.00
54.
811˙
0.00
44.
308˙
0.00
511
.94˙
0.01
51.0˙
1.2
3.84
2˙
0.00
33.
881˙
0.00
64.
848˙
0.01
04.
331˙
0.00
512
.29˙
0.02
85.4˙
2.3
3.75
6˙
0.00
13.
807˙
0.00
24.
745˙
0.00
24.
236˙
0.00
311
.53˙
0.01
52.4˙
1.2
3.83
6˙
0.00
33.
861˙
0.00
74.
851˙
0.01
04.
324˙
0.00
612
.19˙
0.02
100.
8˙
2.8
3.72
6˙
0.00
13.
776˙
0.00
24.
716˙
0.00
24.
213˙
0.00
311
.29˙
0.01
55.7˙
1.3
3.82
9˙
0.00
33.
856˙
0.00
64.
841˙
0.00
54.
318˙
0.00
612
.13˙
0.02
101.
2˙
2.8
3.72
5˙
0.00
13.
771˙
0.00
24.
716˙
0.00
24.
219˙
0.00
411
.29˙
0.01
58.9˙
1.4
3.82
0˙
0.00
33.
848˙
0.00
54.
823˙
0.00
44.
310˙
0.00
512
.04˙
0.01
107.
2˙
3.1
3.71
4˙
0.00
33.
751˙
0.00
34.
706˙
0.00
54.
201˙
0.00
411
.16˙
0.01
71.9˙
1.8
3.78
7˙
0.00
33.
826˙
0.01
24.
766˙
0.01
54.
257˙
0.00
711
.68˙
0.02
107.
6˙
3.1
3.71
3˙
0.00
13.
754˙
0.00
24.
705˙
0.00
24.
201˙
0.00
311
.17˙
0.01
74.6˙
1.9
3.78
0˙
0.00
33.
818˙
0.00
64.
757˙
0.00
64.
259˙
0.00
511
.64˙
0.02
117.
5˙
3.5
3.69
6˙
0.00
13.
747˙
0.00
54.
675˙
0.00
44.
191˙
0.00
511
.05˙
0.01
79.5˙
2.1
3.76
9˙
0.00
23.
807˙
0.00
64.
742˙
0.00
54.
244˙
0.00
511
.53˙
0.02
118.
9˙
3.5
3.69
3˙
0.00
13.
742˙
0.00
24.
680˙
0.00
24.
185˙
0.00
311
.03˙
0.01
84.1˙
2.2
3.75
9˙
0.00
33.
793˙
0.00
54.
730˙
0.00
54.
240˙
0.00
611
.45˙
0.01
123.
4˙
3.7
3.68
7˙
0.00
13.
714˙
0.00
64.
683˙
0.01
04.
183˙
0.00
710
.95˙
0.02
91.8˙
2.7
3.73
3˙
0.00
23.
787˙
0.00
54.
719˙
0.00
44.
213˙
0.00
511
.34˙
0.01
126.
3˙
3.8
3.68
1˙
0.00
13.
725˙
0.00
44.
668˙
0.00
24.
166˙
0.00
310
.90˙
0.01
100.
0˙
2.9
3.72
7˙
0.00
33.
780˙
0.00
54.
696˙
0.00
34.
200˙
0.00
411
.22˙
0.01
129.
5˙
3.9
3.67
6˙
0.00
13.
710˙
0.00
44.
644˙
0.00
24.
138˙
0.00
310
.73˙
0.01
109.
2˙
3.1
3.71
0˙
0.00
23.
749˙
0.00
54.
688˙
0.00
44.
190˙
0.00
511
.09˙
0.01
129.
8˙
3.9
3.67
5˙
0.00
13.
711˙
0.00
24.
642˙
0.00
24.
134˙
0.00
310
.73˙
0.01
111.
0˙
3.2
3.70
6˙
0.00
23.
743˙
0.00
54.
686˙
0.00
44.
187˙
0.00
511
.05˙
0.01
130.
3˙
3.9
3.67
7˙
0.00
13.
704˙
0.00
44.
644˙
0.00
34.
143˙
0.00
310
.73˙
0.01
114.
1˙
3.4
3.69
8˙
0.00
23.
744˙
0.00
54.
677˙
0.00
44.
177˙
0.00
411
.01˙
0.01
117.
0˙
3.4
3.69
7˙
0.00
23.
736˙
0.01
54.
670˙
0.00
74.
167˙
0.00
810
.94˙
0.03
118.
6˙
3.5
3.69
4˙
0.00
23.
721˙
0.00
84.
671˙
0.00
64.
174˙
0.00
610
.92˙
0.02
122.
9˙
3.7
3.68
4˙
0.00
23.
721˙
0.00
64.
653˙
0.00
54.
167˙
0.00
510
.86˙
0.01
126.
0˙
3.8
3.68
2˙
0.00
23.
715˙
0.00
64.
649˙
0.00
54.
161˙
0.00
510
.82˙
0.01
127.
2˙
5b2.
904˙
0.00
4c3.
731˙
0.00
94.
652˙
0.01
24.
161˙
0.01
010
.87˙
0.03
127.
9˙
3.9
3.67
9˙
0.00
23.
720˙
0.00
44.
650˙
0.00
54.
151˙
0.00
510
.80˙
0.01
129.
7˙
3.9
3.67
6˙
0.00
23.
705˙
0.00
74.
633˙
0.00
94.
155˙
0.00
610
.73˙
0.02
132.
3˙
4.1
3.67
2˙
0.00
23.
714˙
0.00
64.
633˙
0.00
54.
133˙
0.00
510
.70˙
0.01
133.
1˙
5b2.
892˙
0.00
3c3.
710˙
0.00
64.
646˙
0.01
04.
136˙
0.00
510
.73˙
0.02
136.
1˙
4.2
3.66
6˙
0.00
23.
699˙
0.00
54.
630˙
0.00
54.
138˙
0.00
510
.66˙
0.01
138.
2˙
5b2.
882˙
0.00
3c3.
704˙
0.01
54.
643˙
0.01
64.
133˙
0.02
310
.70˙
0.05
140.
3˙
4.4
3.66
0˙
0.00
23.
700˙
0.00
54.
622˙
0.00
44.
126˙
0.00
410
.62˙
0.01
142.
1˙
5b2.
875˙
0.00
3c3.
701˙
0.00
64.
633˙
0.01
04.
119˙
0.00
610
.60˙
0.03
145.
7˙
5b2.
869˙
0.00
3c3.
697˙
0.01
04.
623˙
0.01
14.
115˙
0.01
610
.58˙
0.05
149.
3˙
5b2.
863˙
0.00
4c3.
692˙
0.01
54.
615˙
0.01
64.
105˙
0.02
010
.53˙
0.05
a Uni
tcel
lpar
amet
erfo
rpre
ssur
est
anda
rd(g
old
orne
on)i
slis
ted
unde
rap.
b Neo
nlo
aded
.c U
nitc
ellp
aram
eter
ofne
on,a
llot
hera
refr
omgo
ld.
4
Tabl
e4.
Latti
cePa
ram
eter
s(a,
b,c)
and
Uni
tCel
lVol
umes
(V)f
orA
llD
iffra
ctio
nPa
ttern
sfor
the
Stis
hovi
tean
dC
aCl 2
-Typ
ePh
ases
a
Pure
Silic
aSi
lica
With
Alu
min
a
P(G
Pa)
a p(Å
)a(
Å)
b(Å
)c(
Å)
V(m
ol/c
m3 )
P(G
Pa)
a p(Å
)a(
Å)
b(Å
)c(
Å)
V(m
ol/c
m3 )
04.
182˙
0.00
22.
666˙
0.00
214
.04˙
0.01
04.
202˙
0.00
22.
676˙
0.00
214
.22˙
0.01
7˙
0.3
4.02
8˙
0.00
34.
153˙
0.00
22.
652˙
0.00
213
.77˙
0.01
4.0˙
0.1
4.04
8˙0.
002
4.17
7˙
0.00
32.
673˙
0.00
214
.04˙
0.01
18.0˙
0.4
3.96
6˙
0.00
34.
105˙
0.00
32.
641˙
0.00
413
.39˙
0.02
21.4˙
0.4
3.95
0˙
0.00
24.
100˙
0.00
52.
620˙
0.00
513
.30˙
0.03
31.2˙
0.6
3.90
9˙
0.00
34.
048˙
0.00
42.
632˙
0.00
512
.98˙
0.03
38.0˙
0.8
3.88
4˙
0.00
24.
059˙
0.00
42.
591˙
0.00
512
.84˙
0.03
39.2˙
0.8
3.88
0˙
0.00
34.
031˙
0.00
42.
605˙
0.00
312
.74˙
0.02
56.3˙
1.3
3.82
7˙
0.00
24.
075˙
0.01
23.
908˙
0.00
62.
600˙
0.01
112
.46˙
0.03
51.0˙
1.2
3.84
2˙
0.00
34.
024˙
0.00
93.
952˙
0.00
92.
599˙
0.00
512
.44˙
0.02
77.6˙
2.0
3.77
4˙
0.00
44.
031˙
0.00
43.
839˙
0.00
32.
548˙
0.00
211
.87˙
0.01
60.5˙
1.4
3.81
5˙
0.00
34.
022˙
0.00
93.
899˙
0.00
92.
567˙
0.00
512
.12˙
0.02
86.4˙
2.3
3.75
4˙
0.00
44.
018˙
0.00
33.
818˙
0.00
32.
540˙
0.00
111
.73˙
0.01
68.4˙
1.6
3.79
5˙
0.00
34.
023˙
0.00
93.
865˙
0.00
82.
558˙
0.00
511
.98˙
0.02
94.7˙
2.6
3.73
7˙
0.00
34.
004˙
0.00
33.
799˙
0.00
32.
529˙
0.00
111
.58˙
0.01
72.6˙
1.8
3.78
5˙
0.00
34.
015˙
0.00
93.
848˙
0.00
82.
551˙
0.00
511
.86˙
0.02
98.0˙
2.7
3.73
1˙
0.00
33.
997˙
0.00
23.
795˙
0.00
22.
528˙
0.00
111
.54˙
0.01
79.7˙
2.1
3.76
9˙
0.00
34.
005˙
0.00
93.
831˙
0.00
82.
540˙
0.00
511
.73˙
0.02
102.
9˙
2.9
3.72
2˙
0.00
33.
991˙
0.00
33.
782˙
0.00
32.
519˙
0.00
111
.45˙
0.01
84.6˙
2.2
3.75
8˙
0.00
34.
002˙
0.00
93.
819˙
0.00
82.
534˙
0.00
511
.66˙
0.02
107.
2˙
3.0
3.71
4˙
0.00
33.
984˙
0.00
33.
766˙
0.00
32.
512˙
0.00
111
.35˙
0.01
89.2˙
2.4
3.74
8˙
0.00
23.
986˙
0.00
43.
812˙
0.00
52.
528˙
0.00
311
.56˙
0.02
108.
1˙
3.1
3.71
2˙
0.00
33.
970˙
0.00
33.
764˙
0.00
52.
510˙
0.00
211
.29˙
0.02
95.0˙
2.6
3.73
7˙
0.00
23.
981˙
0.00
23.
801˙
0.00
42.
523˙
0.00
211
.50˙
0.01
112.
0˙
3.2
3.70
5˙
0.00
33.
961˙
0.00
43.
758˙
0.00
52.
501˙
0.00
411
.21˙
0.02
117.
5˙
3.5
3.69
6˙
0.00
43.
960˙
0.00
23.
744˙
0.00
22.
504˙
0.00
211
.18˙
0.01
a Uni
tcel
lpar
amet
erof
gold
islis
ted
unde
rap.
Figure 1. Pressure-temperature conditions for the obser-vation of the CaCl2 type (blue symbols) and seifertite (redsymbols) in pure SiO2 (solid symbols), 10 mol % Al2O3-bearing SiO2 (open symbols). In a pure SiO2 sample, wefound a mixture of the CaCl2 type and seifertite (greencircle), which may indicate the P-T conditions correspondto the univariant phase boundary. Laser heating duringdecompression (right triangles) and compression (left tri-angles) is conducted for the Al-bearing samples. The opensquares represent the observations of the CaCl2 type byShieh et al. [2005]. A red solid line is the seifertite phaseboundary estimated from our data. The thin dashed blackline at lower P represents the seifertite formation bound-ary by Dubrovinsky et al. [2001]. The thin black dashedand solid lines at higher P are the CaCl2-type $ seifer-tite boundary proposed by Murakami et al. [2003] andits pressure-corrected ones for direct comparison with ourdata (see text for detail). Blue polygons show the pressure-temperature conditions of the postperovskite boundary inmid-ocean ridge basalt (MORB) (lower and upper boundsfor the thickness of the perovskite-post perovskite mixedphase region (dark and light areas)) [Grocholski et al., 2012].The melting line and the thick gray lines representing theHugoniots of SiO2 using different silica polymorphs as start-ing material are from Akins and Ahrens [2002]. The kinksare due to the crystalline phase changes.
(e.g., Figure 2d) and disappear after unloading the Ar-loadedsamples (Figure 4).
[13] The synthesis of seifertite is not sensitive to laser cou-pler, as we used iron to synthesize the mineral at 152 GPaand 2800 K from a separate pure SiO2 sample (Ne-5) insteadof Au or Pt (Figure 2d). We found an unidentified peakin this sample at slightly higher 2� than the 002 peakfrom hcp iron. This peak is likely from Fe3C and/or Fe3C7[Sata et al., 2010; Nakajima et al., 2011] produced duringthe laser heating at megabar pressures and suggestive ofcarbon contamination.
[14] A third pure SiO2 glass (Ar-3) was heated at 136 GPaand �3100 K where we observed instantaneous crystal-lization of a mixture of seifertite and CaCl2-type structure(Figure 2b). The observation of both low- and high-pressurephases suggest that the phase boundary may exist at those
5
GROCHOLSKI ET AL.: PROPERTIES OF SEIFERTITE
Figure 2. X-ray diffraction patterns of the silica poly-morphs at high pressure-temperature. (a) The CaCl2 type at130 GPa and 2500 K in Al-bearing SiO2 (Ne-6Al). (b) Mix-ture of seifertite and CaCl2 type in Al-free SiO2 synthesizedat 136 GPa and 3100 K. Molecular nitrogen captured duringcryogenic Ar loading (N2?) could be partially responsiblefor the large peak that should be a weak seifertite reflection.(c) Seifertite at 139 GPa and 2500 K in pure SiO2 (Ar-1).(d) Temperature quenched pure SiO2 seifertite synthesizedwith Fe laser coupler (Ne-5) at 152 GPa and 2860 K. Laserheating appears to generate some FexCy contamination, butthe diffraction peaks of seifertite are consistent with samplesAr-1 and Ar-2Al using Au as a laser absorber. We providepeak assignments (seifertite: red dots, CaCl2 type: blue dots,unidentified: asterisks; Re: rhenium gasket, Au: gold inter-nal pressure standard, Ar: argon pressure medium, cg-N:the cubic-gauche phase of nitrogen [Eremets et al., 2004]).Nitrogen diffraction peaks are absent in both the Ne-loadedsample (Figure 2d) and when the samples are decompressedto room pressure (Figures 4a and 4b).
P-T conditions. The fourth pure SiO2 sample (Ne-7) washeated at 102 GPa and 2020 K, producing the diffractionpeaks of the CaCl2-type SiO2.
[15] An Al-bearing SiO2 glass (Ar-2Al) was heated at152 GPa and �1960 K where diffuse but identifiableseifertite peaks appear immediately upon laser coupling,sharpening over 30 min of heating. The sample was thendecompressed to 125 GPa and heated for 50 min at 3300 K,where pressure in the sample chamber increased on quench-ing to 130 GPa and seifertite lines sharpened. The samplewas decompressed again to 125 GPa and heated for 50 minbetween 2100 and 3600 K, generating a small amount ofCaCl2 structured silica (a few spots from the 100% intensity110 peak), but with no noticeable decrease in the intensity ofthe seifertite diffraction peaks (Figure 1 and Table 1).
[16] Weak peaks from a few dots in the diffraction pat-terns are found at 0.327 Å–1 (3.058 Å) and 0.381 Å–1
(2.625 Å) (Figure 3c), but do not match with expecteddiffraction lines for the silica polymorphs proposed by Teter
et al. [1998]. The peaks do not appear to be from any Al2O3polymorphs, including the Rh2O3-II-type structure identifiedby Lin et al. [2004]. We also examined the possibility ofthe CaIrO3-type Al2O3 using the data of Ono et al. [2006].Although the line at line at 0.327 Å–1 is close to the 002line of the CaIrO3-type Al2O3, this reflection is expectedto have very low intensity (4%). The most intense line ofthe CaIrO3-type Al2O3 should exist at 0.412 Å–1 (2.427 Å),which is between the most intense lines of cg-N and seifer-tite in Figure 3c. We conclude that these unidentified peaksare not likely from Al2O3. We cannot rule out the peaksare either from some type of unknown exsolved aluminumsilicate or due to a small modification of the ˛-PbO2 typestructure from oxygen defects introduced by the incorpora-tion of Al [Escudero and Langenhorst, 2012]. We note thatwe do not observe these peaks in pure SiO2.
[17] We also cannot rule out the possibility of a smallamount of Al2O3 exsolved from the Al-bearing SiO2 startingmaterials because the detection limit of diffraction intensityat high pressure is about 5%. Hirose et al. [2005] reportedAl2O3 solubility of 8 mol % in seifertite, which is slightlylower than the content in our Al-bearing SiO2 starting mate-rial (10%). If Hirose et al. [2005] is correct about Al2O3solubility, at least 8 mol % should be in our samples as well.The systematically higher molar volume of our Al-bearingseifertite compared to pure SiO2 seifertite is consistent with
Figure 3. X-ray diffraction patterns of seifertite at highpressure and 300 K. (a) Pure SiO2 at 140 GPa (Ar-1), whichis synthesized at 152 GPa and 2440 K. (b) Pure SiO2 at119 GPa (Ar-1), which was decompressed from 140 GPa.(c) Al-bearing SiO2 at 121 GPa (Ar-2Al), which is synthe-sized at 138 GPa and 3300 K. For comparison, we calculateddiffraction patterns of the sixfold coordinated silica poly-morphs using the first-principles results by Teter et al. [1998]at 120 GPa: ˛-PbO2 type (seifertite), postquartz (3 � 2 typeor P21/c type), NaTiF4 type, and SnO2 type. The symbols forthe peak assignments are the same as those in Figure 2.
6
GROCHOLSKI ET AL.: PROPERTIES OF SEIFERTITE
Figure 4. X-ray diffraction patterns of seifertite at 1 barand 300 K. (a) Pure SiO2 quenched from �152 GPaand 2440 K (Ar-1). (b) Al-bearing SiO2 quenched from�128 GPa and 2590 K (Ar-2Al). For comparison, we cal-culated diffraction patterns of different dense silica poly-morphs observed at 1 bar: seifertite [Dera et al., 2002],postquartz [Haines et al., 2001], and baddeleyite type (peakpositions from El Goresy et al. [2000] and peak intensitiesfrom McCullough and Trueblood [1959]). The symbols forthe peak assignments are the same as those in Figure 2.
at least some aluminum dissolving into the structure, asshown in Figure 5 (see below for discussion).
[18] A second Al-bearing SiO2 glass (Ne-6Al) was heatedat �120 GPa and �2500 K, producing the diffraction pat-terns of the pure CaCl2-type SiO2 (Figure 2a). SampleAr-4Al was heated at �61–75 GPa and 1750–2560 K, allresulting in CaCl2 structured silica.
[19] Figure 1 shows that our data place the CaCl2 type-seifertite boundary very close to the core-mantle boundary(135 GPa) but still within mantle pressures. The data do notprovide sufficient resolution to determine the effect of Al onthe depth of the seifertite phase transition. Most data pointsin Figure 1 are based on the Au pressure scale [Tsuchiya,2003], except for the data at 70–80 GPa (Ar-4Al) and a sin-gle point for pure silica at 152 GPa (Ne-5) in which Pt andNe were used as pressure scales, respectively. The Pt scale[Holmes et al., 1989] may overestimate pressure comparedto the Au scale by a few GPa at 80 GPa and high temperature[Fei et al., 2007]. A data point at 152 GPa in Figure 1 is froman experiment with Fe as a laser coupler, in which we foundpossible carbon contamination of Fe (Figure 2d). Therefore,we estimated the pressure from Ne peak positions measuredafter heating. The measurement after temperature quenchlikely underestimates pressure because of the thermal pres-sure effect [Heinz, 1990; Andrault et al., 1998b]. Since thesedata are far from the phase boundary, these uncertainties arenot important for our stability diagram.
[20] The boundary presented by Dubrovinsky et al.[2001] was obtained by compressing ˛-PbO2-type-likematerial synthesized initially from cristobalite without pres-sure medium (Figure 1). A subsequent study using longer
heating durations with cristobalite starting material andquasi-hydrostatic conditions results in diffraction peaks notconsistent with the ˛-PbO2 phase, frequently producinga disordered ˛-PbO2-type structure instead [Prakapenkaet al., 2004; Dubrovinsky et al., 2004]. The lack of con-firmation of the lower pressure boundary presented byDubrovinsky et al. [2001] in experiments using high tem-perature (>2000 K), quasi-hydrostatic pressure media, andnon-cristobalite (or ˛-PbO2-type like) starting materialshighlight the complex nature of the silica system and thepotential to generate metastable phases [i.e., Tsuchida andYagi, 1990; Prokopenko et al., 2001; Haines et al., 2001].
[21] The boundary reported by Murakami et al. [2003]was measured using the Pt scale [Holmes et al., 1989],which may overestimate pressure with respect to the Auscale [Tsuchiya, 2003] at 2500 K and 100 GPa by 7 GPaaccording to Hirose [2006] or 16 GPa according to Fei et al.
14
13
12
11
140120100806040200
Pressure (GPa)
4.1
4.0
3.9
3.8
3.7
3.6
a (Å)
4.6
4.5
4.4
4.3
4.2
4.1
5.1
5.0
4.9
4.8
4.7
4.6
b (Å)
Figure 5. Unit cell parameters of seifertite (pure SiO2 (redsolid symbols), Al-bearing SiO2 (red open symbols), withan Ar medium (circles), with a Ne medium measured inthis study (squares)) (Tables 3 and 4). The molar volumesof the CaCl2 type measured in this study are also included(pure SiO2 (blue solid squares), Al-bearing SiO2 (blue opensquares)). Solid and dashed lines are fits to the alumina-free and alumina-bearing samples, respectively (Table 2).We also plot unit cell parameters of SiO2 seifertite reportedin previous studies [Tsuchida and Yagi, 1990; Dubrovinskyet al., 1997, 2001, 2004; Murakami et al., 2003]. For thedata points by Murakami et al. [2003], we present both pres-sures estimated from the Pt scale (gray circles) and pressurescorrected for direct comparison with our data (black circles)(see the text for detail). We also include the unit cell param-eters of the seifertite-like phase in MORB (green symbols)[Hirose et al., 2005; Grocholski et al., 2012].
7
GROCHOLSKI ET AL.: PROPERTIES OF SEIFERTITE
[2007]. When the pressure is corrected for a direct compar-ison with our data (solid lines in Figure 1), their boundaryis at significantly lower pressures than our boundary. Unlikeour measurements, Murakami et al. [2003] did not use apressure-transmitting medium or thermal insulation, whichmay result in larger thermal and stress gradients. For exam-ple, studies have found that deviatoric stresses can changethe high pressure behavior of quartz [Kingma et al., 1993;Haines et al., 2001].
[22] With a pressure medium in the diamond-anvil cell,Shieh et al. [2005] reported stability of the CaCl2 type upto pressures 10 GPa higher than the boundary by Murakamiet al. [2003] (�130 GPa) using the same Pt scale (Figure 1).As mentioned above, because the Pt scale may overestimatethe pressure [Fei et al., 2007; Hirose, 2006] with respectto the Au scale we used, the discrepancy between the Ptand Au scales may explain why Shieh et al. [2005] failedto synthesize seifertite at pressures very close to wherewe observed the synthesis. The low-temperature heatingby Shieh et al. [2005] may also have contributed to themetastable persistence of the CaCl2 type, which combinedwith the observations below provide some evidence of thesluggish nature of the transition (Figure 1).
[23] The Clapeyron slope of the seifertite phase transi-tion is poorly constrained from the previous experiments.Although Murakami et al. [2003] presented a positive slope,a negative Clapeyron slope would fit their data equally well.The slope proposed by Dubrovinsky et al. [2001] appearsto be for a boundary of metastable formation of seifertitefar outside the stability field. Our data are an improvementover the previous studies, although we do not have adequatecoverage to quantitatively determine the slope. The kinet-ics of the silica system appear to be strong, evidenced bythe recovery of the ˛-PbO2 structure at room pressure andthe inability to fully convert the ˛-PbO2 to the lower pres-sure stable phase (CaCl2 structure) near the phase boundary.Nevertheless, our data favor a positive Clapeyron slope (pri-marily from the Al-bearing data) which is consistent with thefirst-principles estimations, 5.5–7 MPa/K [Tsuchiya et al.,2004; Oganov et al., 2005; Driver et al., 2010].
3.2. Equation of State[24] Diffraction peaks at 1 bar were indexed by compari-
son with the seifertite found in the Shergotty meteorite [Deraet al., 2002]. Higher pressure and temperature diffractionpatterns retain the same reflections but are shifted to lower dspacings as the density increases. These diffraction patternsare consistent with the ˛-PbO2-type (Pbcn) silica calculatedby Teter et al. [1998] (Figure 3) and allows for unambiguousindexing to determine lattice parameters and volumes. Theunit-cell parameters were determined as a function of pres-sure during decompression (Figure 5 and Tables 3 and 4),using 8–14 reflections from Ar-1 and Ar-2Al, and 5–7 reflec-tions for Ne-5 in which pure seifertite is observed. Theuse of two different pressure transmitting media and lasercouplers allows us to be confident in our seifertite peakassignments. For Ar-3 in which we synthesized a mixture ofseifertite and the CaCl2-type phase, peak overlaps betweenthese two phases make rigorous volume determination ofseifertite challenging and therefore are not included for theequation of state analysis, but the lattice parameters arestill consistent with those determined from pure seifertite
Figure 6. Microphotograph of an Al-free seifertite samplequenched from�152 GPa and�2440 K to room conditions.The sample chamber was collapsed to a small irregular-shaped pocket due to the escape of Ar during decompressionto 1 bar.
samples within uncertainty. Volumes were also measuredfor stishovite and CaCl2-type silica in Ne-6Al and Ne-7,with V0.
[25] During decompression in experiments Ar-1, Ar-2Al,and Ne-5, the diffraction patterns remain the same withoutany new peaks as we cross into the CaCl2 stability field,indicative of seifertite metastability. Our samples appear tobe relatively quasi-hydrostatic during the decompression, asour Debye rings show little or no deviation from circularand different gold reflections give a similar lattice parame-ter. The measured volumes for the two different pure silicasamples are in agreement regardless of the kind of noblegas pressure medium (Ar or Ne). The volume of Al-bearingseifertite remains slightly higher throughout the pressurerange. In experiments Ar-1 and Ar-2Al, we were able todecompress seifertite to 1 bar and recovered the sample(Figure 6). The measured diffraction patterns, comparedwith the calculated patterns of other dense silica polymorphsfound at 1 bar, i.e., baddeleyite type [El Goresy et al., 2000]and postquartz [Haines et al., 2001], show that seifertitesynthesized at �140 GPa still maintains its structure with-out back-converting to other forms (Figure 4). This strongmetastability of seifertite allowed us to calculate 1 bar den-sity (�0) and measure the P-V relations for a wide pressurerange (Figure 5).
[26] The density of pure SiO2 seifertite at 1 bar calculatedfrom V0 measured in this study is the greatest ever recordedfor a silica polymorph (Table 2). Our density of seifertite is�1.5% greater than that of stishovite [Andrault et al., 2003].Our seifertite density is 1.4% higher than seifertite foundin the Shergotty meteorite [Dera et al., 2002], and 2.1%higher than seifertite synthesized from cristobalite startingmaterials [Dubrovinsky et al., 2001]. Our results are virtuallyidentical to the density obtained from quantum Monte Carlocomputations, with our �0 only 0.1% higher than that from
8
GROCHOLSKI ET AL.: PROPERTIES OF SEIFERTITE
Driver et al. [2010] at room pressure and remaining moredense by only 0.2% at the highest pressures of this study(150 GPa). Comparison of room pressure lattice parame-ters differ by only 0.22%, 0.02%, and –0.1% for a, b, andc, respectively [Driver et al., 2010]. Our high-pressure vol-umes (>70 GPa) are also very close to the values from thedensity functional theory calculation [Oganov et al., 2005],although their �0 is higher by 0.6%. Overall, the agree-ment with the most advanced first-principles calculations isremarkable and represents a breakthrough in the ability ofcomputational methods to accurately predict the propertiesof silica at high pressure.
[27] We fit the P-V data using the finite strain method[Birch, 1978] to a second-order Birch-Murnaghan equationof state (Table 2 and Figure 5) due to the trade-off betweenbulk modulus (K0) and the pressure derivative (K00) in the fit-ting [Bell et al., 1987]. We obtained K0 = 322 ˙ 4 GPa forboth pure SiO2 and Al2O3-bearing seifertite, suggesting thatAl2O3 up to 10 mol % does not change the equation of stateof seifertite. Our K0 is in excellent agreement with compu-tational studies [Oganov et al., 2005; Driver et al., 2010](Table 2) but is greater than experimental results on seifer-tite synthesized from cristobalite, even if the difference inK00 is considered (see the fitting for the data at P > 70 GPain Table 2) [Dubrovinsky et al., 2001]. According to ourdata on the CaCl2-type SiO2 using the same experimen-tal setup, the K0 of seifertite is higher by 1.6% than theCaCl2 type. Our K0 for the CaCl2 type is lower than thatof Andrault et al. [2003], with some of the discrepancy dueto their use of the Pt scale [Holmes et al., 1989]. We pre-fer direct comparison of the seifertite K0 with our data onthe CaCl2 type because of internal consistency between thedata sets.
[28] Lattice constants were fit following the modifiedBirch-Murnaghan equation [Xia et al., 1998]. We obtainedK0 = 245 ˙ 5 GPa and K00 = 3.2 ˙ 0.1 for a0 = 4.0731 Å,K0 = 354˙4 GPa and K00 = 3.9˙0.1 for b0 = 5.0263 Å, andK0 = 426˙3 GPa and K00 = 3.7˙0.1 for c0 = 4.4774 Å. Thea axis is significantly more compressible than the other twoaxes, even considering the trade-off between K0 and K00.The a axis is perpendicular to the close packed layer ofoxygen atoms.
[29] Lattice parameters from previous results [Tsuchidaand Yagi, 1990; Dubrovinsky et al., 1997, 2001; Murakamiet al., 2003] have a lot of scatter compared to ourdata. For the lower pressure measurements (<100 GPa),the use of different starting materials and nonhydrostaticstress may produce seifertite in conjunction with othersimilar structures closely related in free energy [Teteret al., 1998; Prokopenko et al., 2001; Prakapenka et al.,2004; Dubrovinsky et al., 2004], making reliable peakidentification challenging.
[30] The volumes reported at P > 100 GPa by Murakamiet al. [2003] are in reasonably close agreement to our dataeven after we correct for the difference in pressure scale[Fei et al., 2007]. The agreement is coincidental, as theaxial parameters are significantly different from our mea-surements, but have offsetting error that leads to consistentvolumes (Figure 5). Murakami et al. [2003] used a smallnumber of broad diffraction peaks (< 5) from nonhydrostaticconditions, with two to three of those peaks having overlapwith the Pt laser absorber.
4. Implications
4.1. Metastability of Seifertite[31] Seifertite is remarkable in that it can be metastably
preserved to 1 bar from 140 GPa over back conversion to itslower pressure forms, such as quartz, stishovite, or densifiedglass. This metastability hysteresis in pressure is an orderof magnitude greater than diamond but lacks the dramaticchange in bonding character such as the sp2-to-sp3 orbitalhybridization for the graphite-to-diamond transition.
[32] The extreme metastability of seifertite is perhapsdue to an energetically unfavorable transition pathway ashas been suggested for other silica analog compounds[Haines and Leger, 1997]. Back conversion from seifertiteto stishovite may require the formation of an intermediatephase such as baddeleyite from consideration of the sym-metry relationships. However, because the baddeleyite-typestructure has higher Si-O coordination state than seifer-tite, the seifertite to baddeleyite-type transition likely hasa large kinetic barrier. Recovery of the ˛-PbO2-type struc-ture demonstrates that some dense high-pressure phases canbe recovered to room conditions by taking advantage oftransitional pathways rather than fundamental changes inbond character.
[33] Conversely, although the pathway is entirelymetastable, the kinetic barrier to form seifertite from cristo-balite is lower than from stishovite [Donadio et al., 2008].The significant difference in density (and perhaps bulkmodulus to a lesser degree) between seifertite synthesizedin its stability field (this study) and through a metastabletransition from cristobalite starting material [Tsuchida andYagi, 1990; Dubrovinsky et al., 2001; Dubrovinskaia et al.,2001] is notable (Table 2). These differences extend to theindividual lattice parameters that are quite different forseifertite sythesized outside its stability field. Nonhydro-static conditions coupled with large activation barriers forcertain starting materials to convert into the stable phase(in this case stishovite or the CaCl2-type structure) maybe the source of these differences. Furthermore, the lowerdensity of natural seifertite compared to what we obtainin our experiments may be important in understanding thepresence of the mineral in the SNC meteorites.
4.2. Seifertite in Meteorites[34] Seifertite has been found in nature in the Shergotty,
Zagami, and NWA 4734 meteorites [Sharp et al., 1999;El Goresy et al., 2008; Miyahara et al., 2013]. Two ofthe meteorites are part of the achondrite SNC family ofMartian origin, while NWA 4734 is of lunar origin. Animportant constraint on the peak shock pressure of thesemeteorites is the lack of evidence for whole-scale melting[Malavergne et al., 2001; El Goresy et al., 2004; Aoudjehaneet al., 2005] and the appearance of small amounts of sei-fertite along with densified glass and stishovite [El Goresyet al., 2008]. The calculated Hugoniots for silica polymorphs[Akins and Ahrens, 2002] combined with our stability fieldmeasurements (Figure 1) would seem to rule out all pre-cursor materials with the exception of stishovite. However,stishovite would require an unreasonably high peak shockpressures (well over 150 GPa, if kinetic issues at lower tem-peratures are considered) that should result in large-scale
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melting, unless some mechanism exists for extreme pressurelocalization during the shock process.
[35] Low-temperature compression of cristobalite is thelowest pressure pathway for generating seifertite [Tsuchidaand Yagi, 1990; Dubrovinsky et al., 2001; Dubrovinskaiaet al., 2001] and has been suggested as the mechanism bywhich it is formed in meteorites [El Goresy et al., 2008].Previous shock wave experiments on cristobalite have beencollected to �30 GPa and produce densified glass on pres-sure release [Gratz et al., 1993]. The production of glassover ˛-PbO2 structured crystals is not surprising as differ-ences in length scales, timescales, and release path make itdifficult to relate laboratory shock experiments directly tolarge impact events [Luo et al., 2003]. While our data do notfavor any particular precursor starting material (other thanruling out glass), it does firmly establish that the seifertite inthese meteorites is far outside the stability field. However, itis not inconsistent with the prevailing idea of a cristobaliteorigin for the mineral [El Goresy et al., 2008].
4.3. Seismic Structures at the Earth’sLowermost Mantle
[36] Homogenized mantle composition is likely undersat-urated in silica [Ringwood, 1982], but chemically distinctmaterials such as oceanic crust are constantly injected intothe mantle at the subduction zones. The heterogeneity maysurvive on a geologic timescale due to extremely sluggishsolid state diffusion at mantle P-T conditions [Allegre andTurcotte, 1986; Holzapfel et al., 2005]. Oceanic crust con-tains silica-oversaturated sediments and basalts, transportingsilica-rich materials into the mantle through subduction.Recent studies have indicated that the basaltic materials maybe transported to the bottom of the mantle and persist thereover long timescales [Hirose et al., 1999; Nakagawa et al.,2010]. In the lowermost mantle, basalt may contain muchless Mg-silicate (perovskite or postperovskite; �30 mol %)than pyrolite (�70 mol %) but contain significant amountof free silica (�40 mol %) while virtually no free silica isexpected in pyrolite [Xu et al., 2008].
[37] Alternatively, some free silica could exist even withina lower mantle of pyrolite composition. Bina [2010] sug-gested the possible existence of free silica even in thepyrolitic mantle as polycrystalline-armored relics with seis-mically detectable scale. Knittle and Jeanloz [1991] andGoarant et al. [1992] suggested that reaction betweenMg-silicate and liquid iron may produce free silica togetherwith iron oxide and iron silicide. The thickness of the reac-tion zone is unknown but should be limited to near thecore-mantle boundary. Therefore, there are at least threepossibilities for the existence of free silica locally in thelowermost mantle.
[38] We found that the CaCl2 type to seifertite phaseboundary is consistent with a positive Clapeyron slope(Figure 1) and occurs at a pressure very close to the core-mantle boundary in the Earth. Our conclusion is based onthe assumption that the gold scale from Tsuchiya [2003] esti-mates the pressure reasonably well (see discussions in Shimet al. [2008] for the uncertainty in pressure scales). Theincrease in density of 1.5% across the boundary is mostlyoffset by a moderate increase of the bulk modulus (0.7%)(Table 2), resulting in either no change or a slight decrease inbulk sound speed by 0.4˙ 1.0% at 130 GPa. Computational
results suggest shear wave speed should decrease acrossthe transition (�1.0–2.0%) at 0 K [Karki et al., 1997b;Tsuchiya, 2011]. Assuming the 0 K calculations are valid atmantle-relevant temperatures, the D00 discontinuity is asso-ciated with a shear wave increase of 2–3% [Lay et al., 1998;Wysession et al., 1998] and frequently occurs as a shallowerdepth, making it unlikely that the seifertite transition can bethe source of the discontinuity.
[39] The perovskite ! postperovskite transition remainsa better explanation for the D00 discontinuity [Murakamiet al., 2004; Oganov and Ono, 2004]. A decrease inbulk sound speed (–0.8 to –2.4%) [Oganov and Ono, 2004;Iitaka et al., 2004; Shim et al., 2008] and an increase inshear velocity [Oganov and Ono, 2004; Iitaka et al., 2004;Wentzcovitch et al., 2006; Tsuchiya and Tsuchiya, 2011]across the postperovskite transition are more consistent withthe D00 discontinuity [Lay et al., 1998; Wysession et al.,1998; Hutko et al., 2008] than the seifertite transition.
[40] A large boundary thickness due to a wide mixedphase region is found in pyrolitic compositions due to thepartitioning behavior of Fe and Al, making a homogenizedmantle composition an unlikely candidate for the disconti-nuity [Andrault et al., 2010; Catalli et al., 2009]. Grocholskiet al. [2012] confirmed this behavior for pyrolite, but showedthat the mineralogy of basaltic and harzburgitic composi-tions can decrease the mixed phase region. One conclusionfrom this study is that the postperovskite transition may onlybe detectable in the regions with particular types of chem-ical heterogeneities (such as subducted oceanic crust). Thisobservation appears to be consistent with the fact that theD00 discontinuity has been more readily documented in theregions which might be related to the deposition of sub-ducted materials in the lowermost mantle [Lay et al., 1998;van der Hilst et al., 2007; He and Wen, 2011].
[41] The reason for discontinuities at greater depths thanthe D00 discontinuity found in some seismic studies is evenless clear. Thomas et al. [2004a] reported a laterally extend-ing discontinuity at a greater depth than the D00 discontinuityin the lowermost mantle beneath Caribbean. They alsodocumented a deep discontinuity in the lowermost mantlebeneath Eurasia [Thomas et al., 2004b]. Later, Hutkoet al. [2008] also documented a deeper discontinuity at50–80 km depths above the core-mantle boundary. Allof these discontinuities have a shear wave decrease withincreasing depth and have frequently been interpreted asconversion from postperovskite to perovskite. However, theseifertite transition may have a shear wave decrease andshould be considered as a candidate for discontinuities foundbelow the D00 discontinuity.
[42] Some deep discontinuities has been explained by twoseparate intersections between the steep lowermost mantlegeotherm and the postperovskite boundary with a positiveClapeyron slope (“double crossing”) in cold regions of themantle [Hernlund et al., 2005]. The perovskite to post-perovskite transition at a shallower depth would create adiscontinuity with a shear wave speed increase, while thereverse transition at a deeper depth would result in a discon-tinuity with a shear wave speed decrease. This hypothesispredicts a negative correlation in the depths between theshallower and deeper discontinuities, resulting in a varia-tion from single discontinuity to double discontinuity alonga lateral decrease in temperature. Hernlund et al. [2005]
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argued that the single discontinuity found beneath Eurasia[Thomas et al., 2004b] and the double discontinuity foundbeneath the Caribbean [Thomas et al., 2004a] are due tothese temperature differences (high for the former and lowfor the latter cases, respectively).
[43] A lens-shaped double-discontinuity structure docu-mented by van der Hilst et al. [2007] in the lowermostmantle beneath the Caribbean is characterized by a shallowdiscontinuity with an increase in shear wave speed and adeeper discontinuity with a decrease in shear wave speed.The region where a transition occurs from the double discon-tinuity to single discontinuity coincides with a temperatureincrease, as inferred from an overall decrease in shearwave speed from the tomography model. The feature hasbeen interpreted as a structure consistent with the “doublecrossing” model.
[44] In a region slightly west of this, Thomas et al. [2004a]reported a laterally extending double discontinuity structurewith a similar velocity structure. However, Thomas et al.[2004a] found little or even positive correlation in the depthsof these two discontinuities (instead of lens shaped). Anotherdouble-discontinuity structure was found beneath Eurasia,but the correlation between the discontinuities is unclear andthe deeper discontinuity has much less topographic varia-tion [Thomas et al., 2004b]. Therefore, the double crossingmodel may not be universal for all double discontinuitystructures in the D00 region.
[45] Topographic variation in general can be complex ifcomposition and/or mineralogy varies laterally. In addition,the much shorter convection timescale in the liquid ironouter core should smooth the lateral temperature variation atthe core side of the core-mantle boundary, allowing for themantle geotherm to converge approaching the core-mantleboundary [Jeanloz and Morris, 1986]. Lateral temperaturevariation near the deeper discontinuity may be less thanthat near the shallower discontinuity, resulting in a weakcorrelation in the depths of those discontinuities.
[46] An interesting possibility for some of the deep dis-continuities can be the CaCl2 type to seifertite transitionassociated with a shear wave speed decrease. However,because of its positive Clapeyron slope, if the seifertitetransition coexists with the postperovskite transition at ashallower depth, it will likely create a double-discontinuitystructure with a positive correlation in depth variations inresponse to lateral variations in temperature.
[47] Basaltic compositions are particularly interesting inthe framework of multiple discontinuity structures as theycontain both perovskite and CaCl2-type silica at lowermantle pressures. Grocholski et al. [2012] showed that thepostperovskite transition occurs at �120 GPa and 2500 K,which is at a shallower depth than the CaCl2 to seifertitetransition (Figure 1). We note that they used the same pres-sure scale as in this study and therefore direct comparison ofthe pressure-temperature conditions of the phase boundariesis not affected by absolute pressure scale issues. The possi-bility exists for basaltic material alone to create the doublediscontinuity structure with the shallower discontinuity bythe postperovskite transition and the deeper discontinuity bythe seifertite transition. However, it is important to recog-nize basalt is chemically complex and to fully consider if ourresults on the Al2O3-SiO2 binary can be directly applied tothis system.
[48] A direct comparison between our results and pre-vious studies on MORB [Hirose et al., 2005; Grocholskiet al., 2012] is possible because of the similar amount ofalumina in the system. However, we find two main dis-crepancies between the data sets. Incorporation of Al2O3does not appear to strongly decrease the depth of the tran-sition, in contrast to previous studies on MORB [Hiroseet al., 2005; Grocholski et al., 2012] that find a structuraltransition in silica produced in conjunction with postper-ovskite at about 110 GPa. In addition, the post-CaCl2-typesilica phase found in MORB appears to have much higherunit cell volume (�2–4%) [Hirose et al., 2005; Grocholskiet al., 2012] than our alumina-bearing samples (Figure 5).The volumes reported for the CaCl2-type agrees well withour measurements and the use of the same Au pressure scaleeliminates technical issues as the source of the discrepancy.Grocholski et al. [2012] used quasi-hydrostatic thermalinsulation medium like in this study, which might explainslightly lower volume in Grocholski et al. [2012] than inHirose et al. [2005].
[49] Considering these two discrepancies and the smallfree-energy differences among sixfold coordinated silicapolymorphs near this pressure range [Teter et al., 1998], wecannot rule out the possibility of the stability of a differentsilica polymorph in MORB instead of seifertite. AlthoughHirose et al. [2005] reported that the post-CaCl2-type phasedoes not contain impurities other than Al, minor amounts ofother cations such as Na close to the detection limit of theanalytic TEM used for the measurements could potentiallystabilize other silica polymorphs with respect to seifertitein MORB.
[50] On the other hand, the diffraction patterns of MORBat pressures over 100 GPa are very complicated due to theexistence of at least four different phases with three of thembeing low symmetry (orthorhombic). Therefore, unambigu-ous identification for the diffraction peaks and subsequentstructural identification and volume measurements of silicaare very challenging for MORB compositions, especiallyin comparison to these experiments on the SiO2-Al2O3binary system.
[51] The ambiguity outlined above is unlikely applicablefor the case of silica produced in the core-mantle reactionzone [Knittle and Jeanloz, 1991; Goarant et al., 1992] orfrom armored silica relics [Bina, 2010] in normal (pyrolitic)mantle composition. The system contains a smaller amountof Al2O3, and Al2O3 may partition preferentially into theperovskite-postperovskite system. Therefore, the seismicinterpretation of our result on pure silica seems to be moreapplicable for these cases.
5. Conclusion[52] We have synthesized the most dense silica calculated
from the smallest molar volume ever experimentally mea-sured at room pressure. The room pressure volume allowsus to determine the equation of state of seifertite, whichhas remarkable agreement with recent computer simulations.Our experiments using quasi-hydrostatic pressure media andlong heating cycles allow us to avoid the generation ofmetastable phases at high pressure and support a positiveClapeyron slope with stability of the mineral at conditionsclose to the core-mantle boundary of the Earth. This stability
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field supports a metastable transition from cristobalite toexplain seifertite found in meteorites as previously sug-gested since no silica polymorph can be shocked to suchhigh pressures without extensive melting. The slight changein bulk sound speed combined with the computationallypredicted decrease in shear wave velocity (if this holds atthe higher temperatures of the mantle) could explain dis-continuities right above the core-mantle boundary, or somedouble discontinuities that are inconsistent with the doublecrossing model.
[53] Acknowledgments. We would like to thank three anonymousreviewers for helpful suggestions that greatly improved the manuscript.This work was performed at GeoSoilEnviroCARS (Sector 13), AdvancedPhoton Source (APS), Argonne National Laboratory. GeoSoilEnviro-CARS is supported by the National Science Foundation–Earth Sciences(EAR-1128799) and Department of Energy–Geosciences (DE-FG02-94ER14466). Use of the Advanced Photon Source was supported by theU. S. Department of Energy, Office of Science, Office of Basic EnergySciences, under contract DE-AC02-06CH11357. This work is supportedby NSF to S.H.S. (EAR1301813) with additional support to B.G. fromthe Peter Buck Fellowship at the National Museum of Natural History,Smithsonian Institution.
ReferencesAkins, J., and T. Ahrens (2002), Dynamic compression of SiO2: A
new interpretation, Geophys. Res. Lett., 29(10), 1394, doi:10.1029/2002GL014806.
Allegre, C. J., and D. L. Turcotte (1986), Implications of a 2-componentmarble-cake mantle, Nature, 323(6084), 123–127.
Andrault, D., G. Fiquet, F. Guyot, and M. Hanfland (1998a), Pressure-induced landau-type transition in stishovite, Science, 282(5384),720–724.
Andrault, D., G. Fiquet, J. P. Itie, P. Richet, P. Gillet, D. Häusermann,and M. Hanfland (1998b), Thermal pressure in the laser-heateddiamond-anvil cell: An X-ray diffraction study, Eur. J. Mineral., 10,931–940.
Andrault, D., R. J. Angel, J. L. Mosenfelder, and T. Le Bihan (2003),Equation of state of stishovite to lower mantle pressures, Am. Mineral.,88(2-3), 301–307.
Andrault, D., M. Munoz, N. Bolfan-Casanova, N. Guignot, J.-P. Perrillat, G.Aquilanti, and S. Pascarelli (2010), Experimental evidence for perovskiteand post-perovskite coexistence throughout the whole D00 region, EarthPlanet. Sci. Lett., 293, 90–96.
Aoudjehane, H., A. Jambon, B. Reynard, and P. Blanc (2005), Silica asa shock index in shergottites: A cathodoluminescence study, Meteorit.Planet. Sc., 40, 967–979.
Bina, C. R. (2010), Scale limits of free-silica scatters in the lower mantle,Phys. Earth Planet. Int., 183, 110–114.
Birch, F. (1978), Finite strain isotherms for single crystal and polycrys-talline NaCl at high-pressures and 300-degree K, J. Geophys. Res., 83,1257–1268.
Bell, P. M., H.-K. Mao, and J. A. Xu (1987), Error analysis of parameter-fitting in equations of state for mantle minerals, in High-PressureResearch in Mineral Physics: A Volume in Honor of Syun-iti Akimoto,Geophys. Monogr. Ser., vol. 39, edited by M. H. Manghnani and Y.Syono, pp. 447–454, Terra Scientific and AGU, Washington D. C.
Boctor, N., C. Alexander, J. Wang, and E. Hauri (2003), The sources ofwater in Martian meteorites: Clues from hydrogen isotopes, Geochim.Cosmochim. Acta, 67(20), 3971–3989.
Catalli, K., S.-H. Shim, and V. Prakapenka (2009), Thickness andClapeyron slope of the post-perovskite transition, Nature, 462, 782–785.
Dera, P., C. Prewitt, N. Boctor, and R. Hemley (2002), Characterization of ahigh-pressure phase of silica from the Martian meteorite Shergotty, Am.Mineral., 87(7), 1018–1023.
Dewaele, A., F. Datchi, P. Loubeyre, and M. Mezouar (2008), Highpressure-high temperature equations of state of neon and diamond, Phys.Rev. B, 77, 094,106.
Donadio, D., R. Martonak, P. Raiteri, and M. Parrinello (2008), Influenceof temperature and anisotropic pressure on the phase transitions in˛-cristobalite, Phys. Rev. Lett., 100, 165,502.
Driver, K. P., R. E. Cohen, Z. Wu, B. Militzer, P. L. Rios, M. D. Towler, R. J.Needs, and J. W. Wilkins (2010), Quantum monte carlo computations ofphase stability, equations of state, and elasticity of high-pressure silica,P. Natl. Acad. Sci. USA, 107, 9519–9524.
Dubrovinskaia, N., L. Dubrovinsky, S. Saxena, F. Tutti, S. Rekhi, and T.Le Bihan (2001), Direct transition from cristobalite to post-stishovite˛-PbO2-like silica phase, Eur. J. Mineral., 13(3), 479–483.
Dubrovinsky, L., S. Saxena, P. Lazor, R. Ahuja, O. Eriksson, J. Wills, andB. Johansson (1997), Experimental and theoretical identification of a newhigh-pressure phase of silica, Nature, 388, 362–365.
Dubrovinsky, L., N. Dubrovinskaia, S. Saxena, F. Tutti, S. Rekhi, T.Le Bihan, G. Shen, and J. Hu (2001), Pressure-induced transformationsof cristobalite, Chem. Phys. Lett., 333(3-4), 264–270.
Dubrovinsky, L., N. Dubrovinskaia, V. Prakapenka, F. Seifert, F.Langenhorst, V. Dmitriev, H. Weber, and T. Le Bihan (2004), A classof new high-pressure silica polymorphs, Phys. Earth Planet. Int., 143,231–240.
El Goresy, A., L. Dubrovinsky, T. Sharp, S. Saxena, and M. Chen(2000), A monoclinic post-stishovite polymorph of silica in the Shergottymeteorite, Science, 288(5471), 1632–1634.
El Goresy, A., L. Dubrovinsky, T. Sharp, and M. Chen (2004),Stishovite and post-stishovite polymorphs of silica in the Shergottymeteorite: Their nature, petrographic settings versus theoretical pre-dictions and relevance to Earth’s mantle, J. Phys. Chem. Solids, 65,1597–1608.
El Goresy, A., P. Dera, T. G. Sharp, C. T. Prewitt, M. Chen, L. Dubrovinsky,B. Wopenka, N. Z. Boctor, and R. J. Hemley (2008), Seifertite, a denseorthorhombic polymorph of silica from the Martian meteorites Shergottyand zagami, Eur. J. Mineral., 20, 523–528.
Eremets, M., A. Gavriliuk, I. Trojan, D. Dzivenko, and R. Boehler (2004),Single-bonded cubic form of nitrogen, Nat. Mater., 3, 558–563.
Escudero, A., and F. Langenhorst (2012), Aluminum incorporation in˛-PbO2 type TiO2 at pressures up to 20 GPa, Phys. Earth Planet. Inter.,190-191, 87–94.
Fei, Y., A. Ricolleau, M. Frank, K. Mibe, G. Shen, and V. Prakapenka(2007), Toward an internally consistent pressure scale, P. Natl. Acad. Sci.,104, 9182–9186.
Goarant, F., F. Guyot, J. Peyronneau, and J. Poirier (1992), High-pressureand high-temperature reactions between silicates and liquid-iron alloys,in the diamond anvil cell, studied by analytical electron-microscopy, J.Geophys. Res., 97(B4), 4477–4487.
Gratz, A., L. Deloach, T. Clough, and W. Nellis (1993), Shock amorphiza-tion of cristobalite, Science, 259(5095), 663–666.
Grocholski, B., K. Catalli, S.-H. Shim, and V. Prakapenka (2012),Mineralogic effects on the post-perovskite boundary, Proc. Nat. Acad.Sci. USA, 109, 2275–2279.
Haines, J., and J. Leger (1997), X-ray diffraction study of the phasetransitions and structural evolution of tin dioxide at high pressure: Rela-tionships between structure types and implications for other rutile-typedioxides, Phys. Rev. B, 55, 11,144–11,154.
Haines, J., J. Leger, and O. Schulte (1996), Pa3 modified fluorite-typestructures in metal dioxides at high pressure, Science, 271(5249),629–631.
Haines, J., J. Leger, F. Gorelli, and M. Hanfland (2001), Crystallinepost-quartz phase in silica at high pressure, Phys. Rev. Lett., 87,155,503.
Hammersley, A. P., S. O. Svensson, M. Hanfland, A. N. Fitch, andD. Häusermann (1996), Two-dimensional detector software: From realdetector to idealised image or two-theta scan, High Pressure Res., 14,235–248.
He, Y., and L. Wen (2011), Seismic velocity structures and detailed featuresof the D00 discontinuity near the core-mantle boundary beneath easternEurasia, Phys. Earth Planet. Int., 189, 176–184.
Heinz, D. L. (1990), Thermal pressure in the laser-heated diamond anvilcell, Geophys. Res. Lett., 17, 1161–1164.
Hernlund, J. W., C. Thomas, and P. J. Tackley (2005), A doubling of thepost-perovskite phase boundary and structure of the Earth’s lowermostmantle, Nature, 434, 882–886.
Hirose, K. (2006), Post-perovskite phase transition and its geophysicalimplications, Rev. Geophys., 44, RG3001, doi:10.1029/2005RG000186.
Hirose, K., Y. Fei, Y. Z. Ma, and H.-K. Mao (1999), The fate of subductedbasaltic crust in the Earth’s lower mantle, Nature, 397, 53–56.
Hirose, K., N. Takafuji, N. Sata, and Y. Ohishi (2005), Phase transition anddensity of subducted MORB crust in the lower mantle, Earth Planet. Sc.Lett., 237, 239–251.
Holland, T. J. B., and S. A. T. Redfern (1997), Unit cell refinement frompowder diffraction data: The use of regression diagnostics, Mineral.Mag., 61, 65–77
Holmes, N., J. Moriarty, G. Gathers, and W. Nellis (1989), The equationof state of platinum to 660 GPa (6.6 Mbar), J. Appl. Phys., 66(7),2962–2967.
Holzapfel, C., D. Rubie, D. Frost, and F. Langenhorst (2005), Fe-Mg inter-diffusion in (Mg,Fe)SiO3 perovskite and lower mantle reequilibration,Science, 309, 1707–1710.
12
GROCHOLSKI ET AL.: PROPERTIES OF SEIFERTITE
Hutko, A. R., T. Lay, J. Revenaugh, and E. J. Garnero (2008), Anticorre-lated seismic velocity anomalies from post-perovskite in the lowermostmantle, Science, 320, 1070–1074.
Iitaka, T., K. Hirose, K. Kawamura, and M. Murakami (2004), The elasticityof the MgSiO3 post-perovskite phase in the Earth’s lowermost mantle,Nature, 430, 442–445.
Jeanloz, R., and S. Morris (1986), Temperature distribution in the crust andmantle, Ann. Rev. Earth Planet. Sci., 14, 377–415.
Karki, B., M. Warren, L. Stixrude, G. Ackland, and J. Crain (1997a), Abinitio studies of high-pressure structural transformations in silica, Phys.Rev. B, 55, 3465–3471.
Karki, B., L. Stixrude, and J. Crain (1997b), Ab initio elasticity ofthree high-pressure polymorphs of silica, Geophys. Res. Lett., 24(24),3269–3272.
Kingma, K., R. Hemley, H. Mao, and D. Veblen (1993), New high-pressuretransformation in alpha-quartz, Phys. Rev. Lett., 70(25), 3927–3930.
Kingma, K. J., R. E. Cohen, R. J. Hemley, and H.-K. Mao (1995), Transfor-mation of stishovite to a denser phase at lower-mantle pressures, Nature,374, 243–245.
Knittle, E., and R. Jeanloz (1991), Earth’s core-mantle boundary—Resultsof experiments at high pressures and temperatures, Science, 251(5000),1438–1443.
Kuwayama, Y., K. Hirose, N. Sata, and Y. Ohishi (2005), The pyrite-typehigh-pressure form of silica, Science, 309, 923–925.
Lay, T., Q. Williams, and E. Garnero (1998), The core-mantle boundarylayer and deep Earth dynamics, Nature, 392(6675), 461–468.
Lin, J.-F., O. Degtyareva, C. T. Prewitt, P. Dera, N. Sata, E. Gregoryanz,H.-K. Mao, and R. J. Hemley (2004), Crystal structure of a high-pressure/high-temperature phase of alumina by in situ X-ray diffraction,Nature Mater., 3, 389–393.
Luo, S., T. Ahrens, and P. Asimow (2003), Polymorphism, superheat-ing, and amorphization of silica upon shock wave loading and release,J. Geophys. Res., 108(B9), 2421, doi:10.1029/2002JB002317.
Malavergne, V., F. Guyot, K. Benzerara, and I. Martinez (2001),Description of new shock-induced phases in the Shergotty, Zagami,Nakhla and Chassigny meteorites, Meteorit. Planet. Sc., 36(10),1297–1305.
McCullough, J. D., and K. N. Trueblood (1959), The crystal structure ofbaddeleyite (monoclinic ZrO2), Acta Crystallogr., 12, 507–511.
Miyahara, M., S. Kaneko, E. Ohtani, T. Sakai, T. Nagase, M. Kayama, H.Nishido, and N. Hirao (2013), Discovery of seifertite in a shocked lunarmeteorite, Nat. Commun., 4, 855–858.
Murakami, M., K. Hirose, S. Ono, and Y. Ohishi (2003), Stability ofCaCl2-type and ˛-PbO2-type SiO2 at high pressure and temperaturedetermined by in-situ X-ray measurements, Geophys. Res. Lett., 30(5),1207, doi:10.1029/2002GL016722.
Murakami, M., K. Hirose, K. Kawamura, N. Sata, and Y. Ohishi (2004),Post-perovskite phase transition in MgSiO3, Science, 304, 855–858.
Nakagawa, T., P. J. Tackley, F. Deschamps, and J. A. D. Connolly(2010), The influence of MORB and harzburgite composition onthermo-chemical mantle convection in a 3-D spherical shell with self-consistently calculated mineral physics, Earth and Plan. Sci. Lett., 296,403–412.
Nakajima, Y., E. Takahashi, N. Sata, Y. Nishihara, K. Hirose, K.-I.Funakoshi, and Y. Ohishi (2011), Thermoelastic property and high-pressure stability of Fe7C3: Implication for iron-carbide in the Earth’score, Am. Mineral., 96, 1158–1165.
Oganov, A. R., and S. Ono (2004), Theoretical and experimental evidencefor a post-perovskite phase of MgSiO3 in Earth’s D00 layer, Nature, 430,445–448.
Oganov, A., M. Gillan, and G. Price (2005), Structural stability of silica athigh pressures and temperatures, Phys. Rev. B, 71(6), 064,104.
Ono, S., K. Hirose, M. Murakami, and M. Isshiki (2002), Post-stishovitephase boundary in SiO2 determined by in situ X-ray observations, EarthPlanet. Sc. Lett., 197(3-4), 187–192.
Ono, S., A. R. Oganov, T. Koyama, and H. Shimizu (2006), Stability andcompressibility of the high-pressure phases of Al2O3 up to 200 GPa:Implications for the electrical conductivity of the base of the lowermantle, Earth Planet. Sc. Lett., 246, 326–335.
Prakapenka, V., L. Dubrovinsky, G. Shen, M. Rivers, S. Sutton, V. Dmitriev,H. Weber, and T. Le Bihan (2003a), ˛-PbO2-type high-pressure poly-morph of GeO2, Phys. Rev. B, 67(13), 132,101–132,104.
Prakapenka, V., G. Shen, and L. Dubrovinsky (2003b), Carbon trans-port in diamond anvil cells, High Temp.-High Press., 35-6(2), 237–249,doi:10.1068/htjr098.
Prakapenka, V., G. Shen, L. Dubrovinsky, M. Rivers, and S. Sutton (2004),High pressure induced phase transformation of SiO2 and GeO2: Differ-ence and similarity, J. Phys. Chem. Solids, 65(8-9), 1537–1545.
Prakapenka, V. B., A. Kubo, A. Kuznetsov, A. Laskin, O. Shkurikhin, P.Dera, M. L. Rivers, and S. R. Sutton (2008), Advanced flat top laserheating system for high pressure research at GSECARS: Application tothe melting behavior of germanium, High Pressure Res., 28, 225–235,International Workshop on Synchrotron High-Pressure Mineral Physicsand Materials Science, Chicago, Ill., DEC 06–07, 2007.
Prokopenko, V., L. Dubrovinsky, V. Dmitriev, and H. Weber (2001), In situcharacterization of phase transitions in cristobalite under high pressureby Raman spectroscopy and X-ray diffraction, J. Alloys Compd., 327,87–95.
Ringwood, A. E. (1982), Phase-transformations and differentiation in sub-ducted lithosphere—Implications for mantle dynamics, basalt petrogen-esis, and crustal evolution, J. Geol., 90(6), 611–643.
Rivers, M., V. B. Prakapenka, A. Kubo, C. Pullins, C. M. Holl, and S. D.Jacobsen (2008), The COMPRES/GSECARS gas-loading system fordiamond anvil cells at the Advanced Photon Source, High. Pressure Res.,28(3), 273–292.
Sata, N., K. Hirose, G. Shen, Y. Nakajima, Y. Ohishi, and N. Hirao (2010),Compression of FeSi, Fe3C, Fe0.95O, and FeS under the core pressuresand implication for light element in the Earth’s core, J. Geophys. Res.,115, B09204, doi:10.1029/2009JB006975.
Sharp, T. G., A. E. Goresy, B. Wopenka, and M. Chen (1999), A post-stishovite SiO2 polymorph in the meteorite Shergotty: Implications forimpact events, Science, 284, 1511–1513.
Shieh, S., T. Duffy, and G. Shen (2005), X-ray diffraction study of phasestability in SiO2 at deep mantle conditions, Earth Planet. Sc. Lett., 235,273–282.
Shim, S.-H., K. Catalli, J. Hustoft, A. Kubo, V. B. Prakapenka, W. A.Caldwell, and M. Kunz (2008), Crystal structure and thermoelastic prop-erties of (Mg0.91Fe0.09)SiO3 postperovskite up to 135 GPa and 2700 K,Proc. Nat. Acad. USA, 105, 7382–7386.
Stöffler, D., R. Ostertag, C. Jammes, G. Pfannschmidt, P. Gupta, S. Simon,J. Papike, and R. Beauchamp (1986), Shock metamorphism and petrog-raphy of the Shergotty achondrite, Geochim. Cosmochim. Acta, 50(6),889–903.
Tangeman, J., B. Phillips, A. Navrotsky, J. Weber, A. Hixson, and T.Key (2001), Vitreous forsterite (Mg2SiO4): Synthesis, structure, andthermochemistry, Geophys. Res. Lett., 28(13), 2517–2520.
Teter, D., R. Hemley, G. Kresse, and J. Hafner (1998), High pressurepolymorphism in silica, Phys. Rev. Lett., 80, 2145–2148.
Thomas, C., E. J. Garnero, and T. Lay (2004a), High-resolution imaging oflowermost mantle structure under the Cocos plate, J. Geophys. Res., 109,B08307, doi:10.1029/2004JB003013.
Thomas, C., J. M. Kendall, and J. Lowman (2004b), Lower-mantle seismicdiscontinuities and the thermal morphology of subducted slabs, Earthand Planet. Sci. Lett., 225, 105–113.
Tsuchida, Y., and T. Yagi (1990), New pressure-induced transformations ofsilica at room-temperature, Nature, 347, 267–269.
Tsuchiya, T. (2003), First-principles prediction of the P-V-T equation ofstate of gold and the 660-km discontinuity in Earth’s mantle, J. Geophys.Res., 108, 2462, doi:10.1029/2003JB002446.
Tsuchiya, T. (2011), Elasticity of subducted basaltic crust at the lower man-tle pressures: Insights on the nature of deep mantle heterogeneity, Phys.Earth Planet. Int., 188, 142–149.
Tsuchiya, T., and J. Tsuchiya (2011), Prediction of a hexagonal SiO2 phaseaffecting stabilities of MgSiO3 and CaSiO3 at multimegabar pressures,
Proc. Nat. Acad. Sci. USA, 108, 1252–1255.Tsuchiya, T., R. Caracas, and J. Tsuchiya (2004), First principles determi-
nation of the phase boundaries of high-pressure polymorphs of silica,Geophys. Res. Lett., 31, L11610, doi:10.1029/2004GL019649.
van der Hilst, R. D., M. V. de Hoop, P. Wang, S.-H. Shim, P. Ma, and L.Tenorio (2007), Seismostratigraphy and thermal structure of Earth’s core-mantle boundary region, Science, 315, 1813–1817.
Wentzcovitch, R. M., T. Tsuchiya, and J. Tsuchiya (2006), MgSiO3 postper-ovskite at D00 conditions, Proc. Nat. Acad. USA, 103, 543–546.
Wysession, M. E., T. Lay, J. Revenaugh, Q. Williams, E. J. Garnero, R.Jeanloz, and L. H. Kellogg (1998), The D00 discontinuity and its impli-cations, in The Core-Mantle Boundary Region, Geodyn. Ser., vol. 28,pp. 273–297, AGU, Washington, D. C.
Xia, X., D. J. Weidner, and H. Zhao (1998), Equation of state of brucite:Single-crystal Brillouin spectroscopy study and polycrystalline pressure-volume-temperature measurement, Am. Mineral., 83, 68–74.
Xu, W., C. Lithgow-Bertelloni, L. Stixrude, and J. Ritsema (2008), Theeffect of bulk composition and temperature on mantle seismic structure,Earth and Planet. Sci. Lett., 275, 70–79.
Zhang, J., B. Li, W. Utsumi, and R. Liebermann (1996), In situ X-ray obser-vations of the coesite stishovite transition: Reversed phase boundary andkinetics, Phys. Chem. Mineral., 23(1), 1–10.
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