Contrib Mineral Petrol (2010) 160:33–43 DOI 10.1007/s00410-009-0463-9
ORIGINAL PAPER
Antigorite equation of state and anomalous softening at 6 GPa: an in situ single-crystal X-ray diffraction study Fabrizio Nestola • Ross J. Angel • Jing Zhao • Carlos J. Garrido • Vicente Lo´pez Sa´nchez-Vizcaı´no Giancarlo Capitani • Marcello Mellini
•
Received: 20 March 2009 / Accepted: 2 November 2009 / Published online: 19 November 2009 Ó Springer-Verlag 2009
Abstract The compressibility of antigorite has been determined up to 8.826(8) GPa, for the first time by single crystal X-ray diffraction in a diamond anvil cell, on a specimen from Cerro del Almirez. Fifteen pressure– volume data, up to 5.910(6) GPa, have been fit by a thirdorder Birch–Murnaghan equation of state, yielding V0 = ˚ 3, KT0 = 62.9(4) GPa, with K0 = 6.1(2). 2,914.07(23) A The compression of antigorite is very anisotropic with axial compressibilities in the ratio 1.11:1.00:3.22 along a, b and c, respectively. The new equation of state leads to an estimation of the upper stability limit of antigorite that is intermediate with respect to existing values, and in better agreement with experiments. At pressures in excess of 6 GPa antigorite displays a significant volume softening that may be relevant for very cold subducting slabs.
This paper is dedicated to the memory of Prof. Volkmar Trommsdorff whom we knew as a gentleman and great scientist.
Keywords Antigorite High pressure Softening X-ray diffraction
Introduction Antigorite, long disregarded as an alteration serpentine mineral of limited interest, started to receive attention after the piston-cylinder experiments by Ulmer and Trommsdorff (1995) emphasized three important features of its phase relations. It has a wide stability field (e.g., up to 720°C at 2 GPa, or 620°C at 5 GPa) and therefore will be widely distributed in subducting slabs. The breakdown reaction of antigorite has a negative dP/dT slope which means that it is stabilised in cooler regions of subducted slabs. The breakdown reaction results in a huge water release of 13 wt% at 150–200 km depth, which has implications for calc-alkaline magma generation and mantle geodynamics. Similar results were obtained by Wunder and Schreyer (1997) for pure MgO–SiO2–H2O (MSH) compositions, while antigorites in the MgO–Al2O3–SiO2–H2O (MASH) system contain
Communicated by M.W. Schmidt. F. Nestola (&) Dipartimento di Geoscienze, Universita` di Padova, Via Giotto 1, 35137 Padova, Italy e-mail:
[email protected]
V. L. Sa´nchez-Vizcaı´no Departamento de Geologı´a, Universidad de Jae´n (Unidad Asociada al CSIC-IACT Granada), Escuela Polite´cnica Superior, Alfonso X El Sabio 28, 23700 Linares, Spain
F. Nestola CNR-IGG, Sezione di Padova, Via Giotto 1, 35137 Padova, Italy
G. Capitani Dipartimento di Scienze Geologiche e Geotecnologie, Universita` degli Studi di Milano Bicocca, P.za della Scienza 4, 20126 Milan, Italy
R. J. Angel J. Zhao Crystallography Laboratory, Virginia Polytechnic Institute and State University, Blacksburg, VA 24061, USA C. J. Garrido Facultad de Ciencias, Instituto Andaluz de Ciencias de la Tierra (IACT), CSIC & UGR, 18002 Granada, Spain
M. Mellini Dipartimento di Scienze della Terra, Universita` di Siena, Via Laterina 8, 53100 Siena, Italy
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aluminium and remain stable to higher temperatures and slightly higher pressures than MSH compositions (Bromiley and Pawley 2003). Following those studies, antigorite serpentinites have become common components in the geological and geophysical modelling of convergent margins (e.g., Tatsumi 1989; Bebout and Barton 1989; Ulmer and Trommsdorff 1995; Sgambelluri et al. 1995; Schmidt and Poli 1998; Guillot et al. 2000, 2001; Peacock 2001; Rupke et al. 2002; Hacker et al. 2003a, b; Tonarini et al. 2007). While exhaustive studies of the defect structures that occur within antigorite have been performed (e.g., Otten 1993; Capitani and Mellini 2005, 2007, and references therein), crystal-chemical and crystal-physical data for antigorite remains limited because of the lack of goodquality single crystals. The first three-dimensional structure refinements were only produced recently (Capitani and Mellini 2004, 2006), while high-temperature and highpressure structural data are not yet available. The compressibility of antigorite was measured by synchrotron Xray powder diffraction (Bose and Navrotsky 1998, Hilairet et al. 2006a, b). Anomalous shifts of the Raman peaks in both natural antigorites (Auzende et al. 2004) and stoichiometric synthetic antigorite (Reynard and Wunder 2006) have been reported at pressures around 6 GPa. Recently, Padro´n-Navarta et al. (2008) described a new occurrence of crystalline, highly-ordered, well-preserved, abundant HP/HT antigorite at Cerro del Almirez in Spain. The large aluminium content makes it suitable for studying serpentinite phase relations within the FeO–MgO–Al2O3– SiO2–H2O (FMASH) system. In particular, the sample has the same Al content as that studied by Hilairet et al. (2006a). Using this Almirez specimen, we now report the first highpressure in situ single-crystal study of antigorite which defines the high-pressure behaviour of antigorite in greater detail than is possible using powder diffraction data. The potential implications for the phase stability and rheology of subducted antigorite serpentinites are discussed.
Experimental methods Sample For a long time the structural analysis of antigorite has been hampered by the absence of adequate single crystals due to their limited size and their containing structure defects, [100] polysomatic faults (001) polysynthetic twins, b/3 stacking faults, etc. (e.g. Otten 1993; Capitani and Mellini 2005, 2007). The first structure refinement was made possible by the discovery of the exceptional Mg159 specimen from Val Malenco (Capitani and Mellini 2004, 2006). A similar highly-ordered antigorite (Padro´n-Navarta et al. 2008) has been found recently in the Cerro del Almirez
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ultramafic massif in Spain (Trommsdorff et al. 1998). A 200 m thick body of metamorphosed antigorite serpentinites are separated through a sharp antigorite-out boundary from their breakdown products of olivine ? enstatite ? chlorite. The Cerro del Almirez antigorite has an Al2O3 content of 3.5–4.0 wt%, which is sufficient to stabilize it at high temperature (Bromiley and Pawley 2003). Pressures of 1.7–2.0 GPa and a temperature of 640°C have been estimated for the antigorite-out dehydration reaction (Lo´pezSanchez Vizcaı´no et al. 2005, 2009) at Cerro del Almirez. Close to the boundary antigorite shows the usual disorder features. But a few meters away from the boundary the antigorite is extremely ordered, with no polysomatic faults over micrometric distances, almost no (001) twins, and no signs of retrograde evolution. It is probable that the high degree of structural order was acquired during long metamorphic annealing at high pressures and temperatures. Since the Almirez rocks are the highest grade serpentinites known, Padro´n-Navarta et al. (2008) suggested that they may be a proxy for the state of serpentinites at depth in subduction zones. For the HP X-ray diffraction study, we selected the Al00-02 specimen of Padro´n-Navarta et al. (2008). The specimen is a foliated serpentinite from the upper part of the Cerro del Almirez ultramafic body. It consists of the ordered m = 17 antigorite polysome, associated with olivine, diopside, magnetite and rare tremolite. Its chemical composition is (Mg2.584Fe0.126Al0.092Cr0.016Ni0.003Mn0.002)R=2.824 (Si1.898Al0.102)R=2.000Cl0.003 (Padro´n-Navarta et al. 2008). High-pressure X-ray diffraction A colourless single crystal of the Almirez Al00-02 antigorite was selected for in situ high-pressure X-ray diffraction measurements. The crystal was free of twins and inclusions, with a size of 70 9 60 9 45 lm3. It was loaded in an ETH-type diamond-anvil cell (DAC) (Miletich et al. 2000) using a T301 steel foil gasket, preindented to 90 lm thickness with a hole 250 lm in diameter. A 4:1 methanol:ethanol mixture was used as pressure-transmitting medium, which remained hydrostatic to the maximum pressures reached (Angel et al. 2007) in this experiment. A quartz crystal was also loaded in the DAC as an internal pressure standard (Angel et al. 1997). The room temperature unit-cell parameters (Table 1) were determined at 19 different pressures up to 8.826 GPa on a Huber four-circle diffractometer (non-monochromatized MoKa radiation) using eight-position centering of at least 20 Bragg reflections at each pressure, according to the procedure of King and Finger (1979). Centring procedures and vector-leastsquares refinement of the unit-cell constants were performed using the SINGLE software (Angel et al. 2000), according to the protocols of Ralph and Finger (1982), without any symmetry constraints in order to monitor any
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Table 1 Unit-cell parameters and unit-cell volumes at different pressures for Almirez antigorite P (GPa)
˚) a (A
˚) b (A
˚) c (A
˚ 3) V (A
a(°)a
b(°)
c(°)*
0.00010
43.521(4)
9.2501(5)
7.2405(4)
2,914.22(33)
89.994(10)
91.203(16)
89.998(6)
0.132(3)
43.500(5)
9.2464(8)
7.2314(5)
2,907.93(44)
89.994(14)
91.206(21)
89.996(8)
0.229(4)
43.487(4)
9.2443(6)
7.2251(4)
2,903.92(37)
89.987(12)
91.207(18)
90.001(7)
0.592(4)
43.432(3)
9.2337(7)
7.2006(4)
2,887.09(34)
89.997(13)
91.175(19)
89.999(7)
1.080(4)
43.361(4)
9.2198(6)
7.1716(4)
2,866.54(39)
89.997(17)
91.117(19)
90.003(7)
1.753(5)
43.261(5)
9.2011(6)
7.1364(4)
2,840.10(39)
89.999(12)
91.066(20)
89.998(8)
2.482(5)
43.149(4)
9.1803(7)
7.1013(4)
2,812.57(38)
89.988(12)
90.994(19)
90.000(8)
3.692(6)
42.967(4)
9.1460(6)
7.0538(6)
2,771.66(39)
89.998(11)
90.889(19)
90.007(7)
4.926(6)
42.775(4)
9.1109(7)
7.0128(4)
2,732.78(39)
89.997(13)
90.784(19)
90.006(8)
4.355(5)
42.860(5)
9.1259(7)
7.0295(6)
2,749.15(43)
89.986(13)
90.840(20)
89.992(8)
2.844(5)b 1.660(5)b
43.094(4) 43.274(4)
9.1690(6) 9.2035(6)
7.0855(5) 7.1405(5)
2,799.29(40) 2,843.35(39)
89.992(12) 89.994(11)
90.963(19) 91.073(19)
90.000(7) 90.001(7)
2.747(5)
43.114(4)
9.1733(5)
7.0915(4)
2,804.26(34)
89.997(11)
90.990(17)
89.997(6)
5.910(6)
42.621(4)
9.0826(6)
6.9838(5)
2,703.27(38)
90.007(11)
90.702(19)
90.007(11)
7.797(8)
42.354(5)
9.0302(6)
6.9238(7)
2,647.93(43)
89.990(12)
90.720(21)
89.998(8)
8.826(8)
42.228(4)
8.9984(6)
6.8882(6)
2,617.14(40)
89.999(11)
90.849(21)
90.000(8)
8.115(7)b
42.315(5)
9.0,207(6)
6.9129(4)
2,638.50(37)
90.000(12)
90.759(20)
89.992(8)
6.469(5)b
42.537(4)
9.0685(5)
6.9688(4)
2,688.07(33)
89.999(10)
90.668(17)
89.998(7)
4.594(6)b
42.822(4)
9.1200(6)
7.0226(5)
2,742.29(39)
89.999(10)
90.806(18)
89.997(7)
a
Not symmetry constrained to monoclinic
b
Data measured during decompression
possible high-pressure phase transformations. Such experimental procedures allow unit-cell parameters to be determined with a high accuracy and precision, sufficient to monitor even very small anomalies in the compressional behaviour (see Angel et al. 2000 for a review).
Results Equation of state Measurements made on compression and decompression confirmed that the cell parameter changes are completely reversible with no pressure hysteresis, in agreement with Auzende et al. (2004), Reynard and Wunder (2006) and Hilairet et al. (2006a). Figure 1 shows the variation of the unit-cell volume against pressure for the Almirez antigorite, along with the previous data collected by powder diffraction by Hilairet et al. (2006a). Our data show that the volume decreases by more than 10% to 8.826 GPa, with significant curvature in the pressure–volume (P–V) trend. In addition, above 6.5 GPa the measured P–V data deviate from the extrapolation of the lower-pressure data (Fig. 1). The anomalous behaviour of antigorite above 6.5 GPa will be discussed in a separate section below. Figure 2 shows the P–V data transformed into a FE–fE plot in which FE is the ‘‘normalized pressure’’ defined as
Fig. 1 Evolution of the unit-cell volume as a function of pressure for Almirez antigorite. The black curve represents the BM3-EoS refined for data up to 5.910 GPa. For purpose of comparison the data by Hilairet et al. (2006a) have been reported as open symbols. The estimated standard deviations are smaller than the symbols used for studies
FE = P/3 9 fE 9 (1 ? 2fE)5/2 and fE is the ‘‘Eulerian strain’’ defined as fE = [(V0/V)2/3 - 1]/2 (details in Angel 2000). Our lower-pressure data fall on a straight line in the FE–fE plot with a positive slope. This indicates that the P–V data up to 5.910 GPa are adequately described by a 3rd
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Fig. 2 FE–fE plot for the unit-cell volume of Almirez antigorite (black symbols) and for antigorite investigated by Hilairet et al. (2006a) (light grey symbols); the normalized pressure FE is defined as FE = P/3 9 fE 9 (1 ? 2fE)5/2 and the Eulerian strain fE as fE = [(V0/ V)2/3 - 1]/2 (see Angel 2000). The square black symbols for Almirez antigorite correspond to the data measured at pressures lower than 5.91 GPa, whereas and the circle black ones those relative to pressure greater than 5.91 GPa. The square light grey symbols for Hilairet et al. (2006a) are relative to the data measured to pressures lower than 6 GPa, whereas the circle light grey symbols are relative to pressures above 6 GPa. We have decided to use two different symbols for the data of Hilairet et al. (2006a) based on the remarkable change of slope in the FE–fE plot
order Birch–Murnaghan EoS with a value of K0 greater than 4. The anomalous compressional behaviour of antigorite above 6 GPa appears as a sharp change of slope in the FE–fE plot (Fig. 2). Such a sharp change shows that no single equation of state can be used to fit all the data throughout the entire pressure range. The 15 P–V data up to 5.910 GPa, which excludes those close to and above the compressional anomaly, were fit using a third-order Birch–Murnaghan equation of state (BM3-EoS, solid line shown in Fig. 1) (Birch 1947) with the software EOSFIT-5.2 (Angel 2000). Simultaneous refinement of the unit-cell volume V0, the bulk modulus KT0 and the first pressure derivative K0 , resulted in the ˚ 3, KT0 = following coefficients: V0 = 2,914.07(23) A 2 62.9(4) GPa, K0 = 6.1(2) with vw = 1.03 and DPmax = 0.025 GPa. A re-analysis of the data published by Hilairet et al. (2006a) and shown in Fig. 2 as light grey symbols clearly shows that their data cannot be fit by a single equation of state due to a clear discontinuity around 6 GPa (above fE about 0.020). Moreover, the data below 6 GPa (fE *0.020) must be fit by a third-order Birch–Murnaghan EoS as indicated by the sloping line on which the data lie (Angel 2000). The resulting bulk modulus, represented by the intercept of the linear weighted regression (dashed light
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grey line in Fig. 2), is K0 = 57(2) GPa. However, Hilairet et al. (2006a) failed to recognise both the discontinuity and the slope in the FE–fE plot. The value of K0 of 62 or 67 GPa they reported does not represent their data, so the apparent agreement with our value of K0 = 62.9(4) GPa is merely fortuitous. Moreover, our bulk modulus is also significantly stiffer than that obtained on an Al-poor natural specimen by synchrotron powder diffraction by Bose and Navrotsky (1998), who reported KT0 = 49.6(7) GPa with a first pressure derivative K0 = 6.14(43). Hilairet et al. (2006a) suggested that the data of Bose and Navrotsky (1998) may be affected by deviatoric stress at low and medium temperatures due to the use of a DIA-type apparatus with a solid pressure medium which may have also induced errors with pressure calibration. The use of non-hydrostatic solid pressure medium will certainly affect the volume evolution of a powder sample, but the exact effect depends on the elastic anisotropy of the material and its preferred orientation, as well as the same factors for the pressure markers. Axial anisotropy When plotted as relative compression (Fig. 3a), the evolution of the unit-cell parameters with pressure appears very anisotropic, with the c axis significantly more compressible than the other two axes. While b decreases approximately linearly up to the maximum pressure of 8.826 GPa, a and c shows clear curvature up to 5.910 GPa and a change in slope, marked by increased compressibility, at higher pressures, as already noted for the unit-cell volume. Figure 3b shows that the monoclinic b unit-cell angle decreases linearly up to 5.910 GPa, and then increases up to 8.826 GPa. Therefore, the data indicate that the anomalous behaviour shown by the unit-cell volume is strongly coupled with both the c axis and the b angle. The anomalous behaviour of the c axis is also evident in Fig. 4, in which FE–fE plots have been calculated for the three axes. There is a strong change in slope of the plot of the c unit-cell parameter, indicating that it must be fit with a 4th order EoS, whereas no slope changes occur for the a and b axes. The anisotropy of compression was quantified using a parameterised form of the BM3-EoS in which the individual axes are cubed and fit as volumes (Angel 2000). The EoS coefficients are reported in Table 2, being obtained, as for the unit-cell volume, from the 15 data up to 5.910 GPa. The axial compressibilities, calculated using the relation for unit-cell parameters of baxis = -1/3KT0axis (Angel 2000), are also reported in Table 2; this relation, although not commonly used, is the only one that correctly takes into account the real high-pressure behaviour of a single axis direction. The ratio of the axial compressibilities for our antigorite at room pressure is 1.11:1.00:3.22, with the c axis
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Fig. 4 FE–fE plots (Angel 2000), for a, b and c directions of Almirez antigorite. The open symbols represent the data measured at pressures greater than 5.910 GPa Table 2 Refined EoS coefficients and refined unit-cell parameters, obtained by a third-order (a and b parameters and unit-cell volume) and fourth-order (c parameter) Birch–Murnaghan fit for Almirez antigorite
Fig. 3 a Evolution of the unit-cell parameters as a function of pressure, shown as relative compression for Almirez antigorite. The open symbols represent the data measured at pressures greater than 5.910 GPa. b Evolution of the b angle with pressure for Almirez antigorite (black symbols) and for antigorite investigated by Hilairet et al. (2006a) (light grey symbols). The black circle symbols for Almirez antigorite represent the data measured at pressures greater than 5.910 GPa
much more compressible than a and b, which compress approximately at the same rate. For sake of comparison with previous data, linear compressibility can be also calculated as bl = l-1 0 (ql/qP), giving (between room pressure and 5.910 GPa): ba = -3.50(5) 9 10-3 (GPa-1), bb = -3 -1 -3.06(4) 9 10 (GPa ), and bc = -6.0(2) 9 10-3 (GPa-1), with a compressibility anisotropy scheme 1.14:1.00:1.96. These ratios are significantly different from those calculated using the correct b = -1/3KT0axis relation (Table 2). The difference is mainly due to the strong curvature of the evolution of the c axis with pressure, which is not taken into account in the calculation of the linear compressibility.
˚) a0 (A
43.521(2)
Ka0 (GPa)
98(1)
K0
-0.7(3)
b (GPa-1) ˚) b0 (A
-0.00340(5) 9.2504(3)
Kb0 (GPa)
109(1)
K0 b (GPa-1) ˚) c0 (A
-0.1(3) -0.00306(3)
Kc0 (GPa)
33.8(7)
K0
7.0(1.0)
00
7.2406(4)
K
1.3(7)
b (GPa-1) ˚ 3) V0 (A
-0.0099(2) 2,913.99(30)
KT0 (GPa)
62.9(4)
K0
6.1(2) -1
b (GPa )
-0.01590
Anomalous compression The pressure–volume data for the Almirez antigorite shows an unexpected and significant softening above 6 GPa (Fig. 1) that was not recognized in previous studies (e.g., Hilairet et al. 2006a, b). However, the FE–fE plot (Fig. 2) recalculated for the unit-cell volume data reported in Hilairet et al. (2006a), notwithstanding the significant scatter, does show a slope change at fE = 0.025–0.030, exactly in the same pressure range where the anomaly occurs for the
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Fig. 5 FE–fE plot for c direction of antigorite studied by Hilairet et al. (2006a)
Almirez antigorite. Also the FE–fE plot for their c axis data (Fig. 5) shows a change in slope at fE = 0.040, again like the Almirez sample. Finally (Fig. 3b), the b angle in Hilairet et al. (2006a) decreases linearly up to 6–6.5 GPa but then clearly increases, exactly as we have observed from single-crystal data. It therefore appears that the antigorite of Hilairet et al. (2006a) behaves in the same way as the Almirez antigorite sample that we have measured, as would be expected given that the two samples have similar chemical compositions. This softening is presumably the cause of the anomalous shifts in the frequencies of the OH bands in the Raman spectra of synthetic antigorite (Fig. 3 of Reynard and Wunder 2006), instead of being artefacts of the fitting procedure, as suggested by the authors. The variation with pressure of the 3,619 cm-1 Raman mode of Al-rich natural antigorite (Fig. 3b of Auzende et al. 2004) also shows a slope change in the 6–7 GPa pressure range. Both singlecrystal and powder X-ray diffraction as well as Raman spectroscopy therefore indicate a change in structural compression mechanisms of both natural and synthetic antigorite specimens at pressures just above 6 GPa. This would mean that antigorites, regardless the origin, chemical composition and polysome, show significant softening in the pressure interval between 6 and 7 GPa. Recently, Capitani et al. (2009) investigated the crystal structure of the m = 17 polysome of antigorite by first principles methods up to 30 GPa. The authors found changes in compression behaviour, at 3 and 17 GPa, for not thermally corrected calculations. The comparison with the discontinuity experimentally found at 6 GPa is not straightforward. This not only for the differences in pressures at which the experiments and the theoretical
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calculations are performed, but also for the different behaviour shown by b angle, which above 6 GPa turns to increase its value in the experiment and instead decreases even more in the theoretical calculations. Interestingly, Mookherjee and Stixrude (2009) have recently predicted a pressure induced transformation in serpentine lizardite at 7 GPa using molecular dynamic (MD) calculations. Lizardite can be considered to be the minimal (m = 1) polysome of antigorite and this transformation could provide a first order interpretation of the changes in the more complex antigorite structure. The MD calculations indicate that the transformation arises from the differential compressibility of the tetrahedral and octahedral layers that vanishes at about 7 GPa, causing non-linear pressure dependence of some structural and elastic parameters above this pressure. In particular, the C33 (and C11 in less extent) elastic moduli, that are directly related to the compressibilities of the c- and a-axis respectively, decrease with increasing pressure, in good agreement with the results of the present compressional experiment. The transformation is also accompanied by changes in the OH bonding above 7 GPa, that may explain the anomalous shift in the OH frequencies observed in Raman studies of both natural and synthetic antigorites (Auzende et al. 2004; Reynard and Wunder 2006). Similar compression anomalies like that found in antigorite have been reported for different minerals, and have been attributed to phase transitions (e.g. Dera et al. 2003 on dickite Al2Si2O5(OH)4; Nestola et al. 2004 on a Ca–Sr feldspar) or to changes in the compressional deformation mechanism (e.g. Benusa et al. 2005). We do not have any structural data available to determine whether there is a symmetry change in antigorite. But, if the softening that we observed is related to a phase transition, the latter cannot be strongly first order in character as it is not accompanied by a detectable volume discontinuity. However, at pressures above the transition, the difference in volume between that extrapolated from the low-pressure EoS and the actual volume is significant, reaching 0.5% of the total volume, or some 5% of the compression, at 8.826(8) GPa, the maximum pressure reached in this work.
Discussion Implications for antigorite stability in subduction zones Antigorite is the dominant hydroxyl-bearing phase in highpressure hydrated peridotites and the main water carrier in subduction zones at pressures below ca. 5.5 GPa (e.g., Schmidt and Poli 1998 and reference therein). Antigorite dehydration reactions play an important role in the localization of seismic activity and are thought to be the cause of
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the double seismic zones in subduction settings (Peacock 2001; Seno et al. 2001; Dobson et al. 2002). Assessing the importance of serpentinites on the dynamics and seismicity of subduction zones requires knowledge of the phase relations, velocities of seismic waves and density variations of antigorite, all of which rely on precise and accurate elastic properties and EoS for this phase. Inaccurate elastic parameters in thermodynamic databases (e.g., Holland and Powell 1998) and/or an incorrect EoS for antigorite may result into substantial errors in the computation of theoretical phase diagrams (Hilairet et al. 2006a). Figure 6a shows the P–T phase diagram in the MSH system—a simplified chemical system for hydrated peridotite—computed for antigorite with a BM2-EoS from Hilairet et al. (2006a) (dark gray lines; Fig. 6a), with the BM3-EoS from our study (black lines; Fig. 6a), and with the original data of Holland and Powell 1998 (light gray lines crossing at invariant points 1 and 2 of Fig. 6a). As in Hilairet et al. (2006a), other thermodynamic parameters for antigorite (e.g. thermal expansion terms) and for other phases are from the internally-consistent, thermodynamic database of Holland and Powell (1998). Our phase diagram computed with the data of Hilairet et al. (2006a) (Fig. 6a) differs from that calculated by those authors (cf. Fig. 6a of our work and their Fig. 5) because they only modified the value of antigorite KT0 from the database value. By contrast, we have changed both the antigorite KT0 and V0 to the values reported by
Hilairet et al. (2006a) to retain the internal consistency of elastic parameters for antigorite (V0, KT0 and K0 ). The larger antigorite V0 (2,926.23 A3, or 176.22 J bar-1 for an antigorite z = 1) of Hilairet et al. (2006a) relative to that for antigorite in the Holland and Powell (1998) database (V0 = 175.48 J bar-1) results in a reduction in pressure and temperature to which antigorite is stable (compare our Fig. 6a with their Fig. 5). Our antigorite V0 (2,913.99 A3 or 175.48 J bar-1 for an antigorite z = 1) fit to a BM3-EoS for the Almirez highly-ordered antigorite is identical to that in the Holland and Powell (1998) database, which was based upon volume values estimated for Val Malenco antigorites (Mellini et al. 1987). Recalculation of the phase relations with our EoS parameters then shifts the stability of antigorite toward pressures and temperatures intermediate between those derived from the Holland and Powell (1998) database and Hilairet et al. (2006a). Use of the lower bulk modulus value reported by Bose and Navrotsky (1998) into the Holland and Powell (1998) database leads to an overestimate of the stability field of antigorite, towards higher pressures and temperatures. Further assessment of our antigorite EoS data requires comparison with experimentally-determined antigorite equilibria. The two key reactions involving antigorite that are relevant for subduction settings are: Atg = Fo ? En ? H2O and Atg = PhA ? En ? H2O (Fig. 6a, b; mineral abbreviations after Kretz 1983). At intermediate
Fig. 6 a Computed phase relationships in the P–T projections for the MgO–SiO2–H2O system (MSH) using different sets of thermodynamic parameters (KT0, K0 and V0) and equations of state for antigorite: this work (black lines), Hilairet et al. (2006a) (dark grey lines) and Holland and Powell (1998) (light grey lines, only reactions cutting invariant points 1 and 2 are shown). See text for details. Reactions 1–5 are: 1 atg = H2O ? ta ? fo, 2 ta ? fo = H2O ? en, 3 atg = H2O ? ta ? phA, 4 ta ? phA = H2O ? en, and 5 ta ? phA = atg ? en. Reaction equations are written with the high temperature assemblage to the right of the = sign. Mineral names
abbreviations as in Kretz (1983). b. Comparison of the P–T conditions for the two main antigorite breakdown reactions (black solid lines) computed in this work (see a) and those experimentally obtained from Al2O3-free antigorites: B&P (Bromiley and Pawley 2003; the grey horizontal bars represent the width of the brackets); B&P (T) (Bromiley and Pawley 2003 after a correction of temperature as explained by Komabayashi et al. (2005); and K (Komabayashi et al. 2005). Note that our results fit the experimental data by Bromiley and Pawley at P \ 4 GPa and those of Komabayashi et al. (2005) at P [ 4 GPa
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to high pressures, reactions involving talc and antigorite may also be relevant for silica-rich serpentinite compositions (Pawley 1998; Manning 1995; Padro´n-Navarta et al. 2009), but they will not be considered here. Although experimental results are ambiguous (Komabayashi et al. 2005; and references therein), it seems that other components, notably Al2O3 (Bromiley and Pawley 2003) and, to a lesser extent, FeO (Ulmer and Trommsdorff 1995), may expand the stability of antigorite toward higher temperatures and pressures. However, the lack of thermodynamic solution models including the Al2O3-rich antigorite endmember precludes precise thermodynamic modeling of antigorite reactions in the MASH and FMASH systems. Furthermore, it is unlikely that compositional variations (quite limited as far as natural antigorites are concerned) have a great impact on the compressibility of antigorite, compared to those due to polysomatism, pressure and temperature. Therefore, comparison to experimental results in the MSH system at different pressures seems appropriate for the assessment of our antigorite elastic parameters and the suitability of a BM3-EoS for modeling antigorite equilibria in natural systems. The Atg = Fo ? En ? H2O and Atg = PhA ? En ? H2O reactions computed with an antigorite BM3-EoS with our V0, KT0 and K0 match those determined experimentally in system MSH, especially at intermediate and high pressures, where correct modeling of antigorite compressibility is crucial (Fig. 6b). At low and intermediate pressures below the MSH invariant point Atg ? En ? Fo ? phA ? H2O, our calculated Atg = Fo ? En ? H2O equilibrium lies within the P–T range for this reaction reported by Bromiley and Pawley (2003) for Al2O3-free antigorite (Fig. 6a, grey horizontal bars). Our result also matches the Atg = Fo ? En ? H2O equilibrium determined by Bromiley and Pawley (2003) (Fig. 6b, curve ‘‘B&P’’), especially after correction for temperature calibration issues as discussed in Komabayashi et al. (2005) [Fig. 6b, curve ‘‘B&P (T)’’]. At higher pressures, our results are better gauged against more precise, in situ high-pressure X-ray determinations of the invariant point Atg ? En ? Fo ? phA ? H2O, and reaction Atg = PhA ? Fo ? H2O (Komabayashi et al. 2005). Our phase diagram exactly matches the experimentally determined P–T conditions for this invariant point (5.1 GPa and 550°C; Komabayashi et al. 2005), whereas the use of the EoS of Hilairet et al. (2006a) places the invariant point 0.25 GPa and 20°C too low. Similar good agreement is found for the reaction Atg = PhA ? Fo ? H2O (Komabayashi et al. 2005), at high pressure (Fig. 6b) when we use our new EoS parameters. Despite the uncertainties in the EoS and elastic parameters of H2O and PhA, the excellent match of our computed phase diagrams with experimental data over 6 GPa indicates that the present BM3-EoS is the most accurate EoS for antigorite,
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and confirms that these elastic parameters determined for the Almirez highly ordered antigorite (m = 17) are applicable to antigorites in general. Strain softening of antigorite in subduction settings Our study demonstrates significant softening of antigorite above about 6 GPa (Fig. 1). As the (dP/dT)P,T of a reaction is equal to (DrS/DrV)P,T (i.e., the Clayperon relation), this softening may induce variations in the dP/dT of the reactions above 6.5 GPa through variations of the DrV due to the smaller VP,T of antigorite relative to that predicted by the BM3-EoS extrapolated from lower pressures. This effect would, of course, be more obvious above ca. 8 GPa, where the predicted and measured antigorite volume strongly differ (Fig. 1). To assess the potential effects of antigorite softening in subduction zones, we show in Fig. 7 the P–T phase equilibria in the system MSH compared with the theoretical geothermal gradient in the upper 8 km of the lithospheric mantle of a subducting slabs in cold (Tohoku, Japan), intermediate and hot (Nankai and Costa Rica), and very hot (Cascadia) subduction zones (after Peacock 2001; modified by Hacker et al. 2003a, b). In the MSH system, antigorite is not stable at P [ ca. 7 GPa for temperatures greater than 350°C (Fig. 7), suggesting that antigorite
Fig. 7 Phase diagram P–T projection (Fig. 6) with the PT paths (dark grey shading) of upper 8 km of subducted mantle in Tohoku, Nankai, Costa Rica and Cascadia subduction zones (taken from Hacker et al. 2003a, b). Phase relationships at P [ 6 GPa (light grey shading) should be affected by antigorite softening reported in this paper. This effect, however, would only take place in very cold subduction zones such as that represented by the Tohoku PT path
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softening may have little implications for intermediate to hot subduction zones. In intermediate T subduction zones, dehydration of antigorite occurs through the reaction Atg = En ? Fo ? H2O. The upper pressure limit of this reaction in the system MSH is set by the invariant point where Atg ? En ? Fo ? phA ? H2O coexist. This point occurs at 5.1 GPa and 550°C (Komabayashi et al. 2005), well below the pressure where we observed strain softening in antigorite at room temperature. In the MASH system this reaction is shifted to higher temperatures, but not to higher pressures (Pawley 2003; Bromiley and Pawley 2003; Komabayashi et al. 2005). The dP/dT slope of the antigorite equilibria might be only relevant for very cold subduction settings (\550°C) such as Tohoku (Fig. 7), where antigorite breaks down to PhA ? Fo ? H2O (Figs. 6, 7) at pressures above the invariant point 1 at 5.1 GPa in the MSH system. Yet, the gentle negative slope of this reaction restrains antigorite stability to pressures below 7 GPa, where DV variations due to antigorite softening are too small to substantially modify the Clapeyron slope of the reaction. Softening might, however, be relevant if antigorite remains metastable in the slab to greater depths, or if the softening pressure decreases with increasing temperature. Available experimental data indicate, however, that antigorite-breakdown reactions are characterized by a very fast kinetics (Padro´n-Navarta et al. 2009; Perrillat et al. 2005; Komabayashi et al. 2005) making it unlikely for antigorite to be metastable in subduction settings, but the temperature dependence of the softening pressure remains to be determined. Negative buoyancy and slab subduction In spite of the thermodynamic analysis of very simple systems given above, the experimental uncertainties in the dehydration experiments (see discussion in Komabayashi et al. 2005) cannot exclude the possibility that antigorite may be present in subduction zones at pressures in excess of 6 GPa, or the softening may occur at lower pressures at high temperatures. If so, the softening may have important bearing on the mechanical instabilities, non-linear processes, slab dynamics and generation of seismic waves in cold subducting slabs, further amplifying the general mechanisms summarised by Hilairet et al. (2007). In particular, the increase in compressibility may correspond to non-linear effects, triggered by an endogenous cause (i.e., the change in compressibility pattern). Basically, the possible presence of non-linear effects would mean that, for the same driving force and the same deformation path (e.g., subduction compression), an instability is introduced. As known from theory (e.g., McClung 1979), strain softening instabilities are associated with decreasing deformation
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resistance. The effect of these instabilities has been thoroughly analyzed for systems as different as long thin tensile specimens, plastic crystal deformation, or snow avalanches. In terms of the high-pressure behaviour of antigorite, the mechanical instability may correspond to an abrupt increase of negative buoyancy, determined by the increased compressibility. Thus, even if the antigorite is not undergoing any breakdown or phase transformation, an antigorite serpentinite slab will become anomalously denser due to the softening. Therefore, the slab will contract and accelerate while, at the same time, the almost discontinuous increase in volume variation may create instability in the slab which could lead to triggering of seismic waves.
Conclusions Compressibility data were measured for the first time on a single crystal of antigorite by X-ray diffraction. The availability of high-quality single-crystal samples and the experimental procedure used allowed us to determine the EoS with greater accuracy and precision than previous studies. The use of the new EoS data predicts that the stability field of antigorite is significantly extended towards greater pressures, about 0.6–0.7 GPa, and then toward greater depths in the mantle (e.g., 18–21 km) at a temperature of 350°C than previously calculated (Hilairet et al. 2006a), but in agreement with recent phase equilibria studies (e.g. Komabayashi et al. 2005). Above 6 GPa antigorite shows an anomalous softening never reported before. Although such anomalous compression does not seem to affect the stability field of antigorite, it should be taken into account in evaluating the mechanical properties of subducting slabs in high-pressure, low-temperature regimes. Acknowledgments Marı´a Teresa Go´mez Pugnaire and Jose´ A. Padro´n-Navarta are thanked for providing the sample, unpublished data and helpful comments essential for our understanding of the Almirez HP antigorite. We also thank Dr. James A. Connolly for his help with Perple-X computations and particularly with the implementation of the third-order Birch-Murnaghan equation of state for antigorite. M. Walter and an anonymous referee are thanked for reviewing the work and strongly improving it. We want to thank M. Schmidt for handling the manuscript and for several useful suggestions. CJG and VLSV acknowledge the financial support of the Spanish ‘‘Ministerio de Ciencia e Innovacion (MICINN)’’ research grants CGL2006-04440, CGL2007-61205/BTE and ACI2006-A90580, of the Spanish Council for Research (CSIC) grant 2008-30I014, and of ‘‘Junta de Andalucia’’ research groups RNM-145 and RNM131. MM acknowledges the ‘‘Junta de Andalucia’’ for granting a short-term visiting fellowship to the University of Granada. The highpressure measurements were performed at the Virginia Tech Crystallography Laboratory with the support of NSF grant EAR -0738692.
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