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Effect of iron content on electrical conductivity of ringwoodite, with implications for electrical structure in the transition zone
Physics of the Earth and Planetary Interiors 174 (2009) 3–9
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Effect of iron content on electrical conductivity of ringwoodite, with implications for electrical structure in the transition zone Takashi Yoshino ∗ , Tomoo Katsura Institute for Study of the Earth’s Interior, Okayama University, Misasa, Tottori 682-0193, Japan
a r t i c l e
i n f o
Article history: Received 2 October 2007 Received in revised form 26 August 2008 Accepted 21 September 2008 Keywords: Electrical conductivity Iron content Ringwoodite Transition zone Upper mantle
a b s t r a c t Electrical conductivity of ringwoodite with various iron contents [Fe/(Fe + Mg) = 0.09, 0.2 and 0.3] was measured at pressure (20 GPa) and temperature (up to 1900 K) conditions of the lower part of the mantle transition zone in a Kawai-type multi-anvil apparatus. The conductivity increased with increasing total iron content. All electrical conductivity data were fitted to the formula of electrical conductivity = 0 XFe exp(−H/kT), where 0 is the pre-exponential term, XFe is the mole fraction of iron content in the Mg site, H is the activation enthalpy, k is the Boltzmann constant and T is absolute temperature. The activation enthalpy becomes higher at a certain temperature. At high temperatures, the activation enthalpy decreased from 1.44 to 0.92 eV with increasing total Fe content. At low temperatures less than 1000 K, the activation enthalpy also decreases from 1.15 to 0.74 eV with total Fe content. Dependence of the activation enthalpy on Fe content suggests that the dominant mechanism of charge transport is Fe2+ –Fe3+ hopping (small polaron). Recent electrical conductivity-depth profiles of the transition zone beneath the Pacific Ocean obtained from the electromagnetic induction study shows that the conductivity values between 520 and 660 km depths may be explained by ringwoodite with Fe/(Fe + Mg) = 0.10. On the other hand, assuming a normal geotherm, conductivity values beneath the continent or stable craton are considerably lower than those of ringwoodite with Fe/(Fe + Mg) = 0.10. Taking into consideration results from the global seismic tomographic studies, relatively low conductivity in these regions can be explained by the existence of a cooler region compared with the surrounding mantle, rather than the presence of iron-poor ringwoodite, or a combination of both. © 2008 Elsevier B.V. All rights reserved.
1. Introduction Electrical conductivity can be used to quantify temperature, water content, mineralogy, chemical composition and oxygen fugacity of the Earth’s deep mantle. In situ laboratory measurements of electrical conductivity of the major mantle phases under high-pressure and high-temperature conditions in conjunction with conductivity-depth profiles of the mantle derived from electromagnetic data provide some constraints on the temperature and chemistry in the Earth’s interior (e.g., Katsura et al., 1998; Xu et al., 1998; Yoshino et al., 2006, 2008a). For the iron-bearing silicate minerals under high-temperature conditions corresponding to the Earth’s deep interior, a small polaron conduction (hopping of electron holes between Fe2+ and Fe3+ ) has been considered as a dominant conduction mechanism (e.g., Schock et al., 1989; Hirsch et al., 1993; Katsura et al., 1998; Xu et al., 1998; Yoshino et al., 2006, 2008a). For this mechanism, the
∗ Corresponding author. Tel.: +81 858 34 3737; fax: +81 858 34 3450. E-mail address: [email protected] (T. Yoshino). 0031-9201/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.pepi.2008.09.015
conductivity of iron-bearing silicate minerals should be strongly controlled by the total Fe content and Fe3+ / Fe (i.e., oxidation condition). Thus, knowledge of electrical properties of silicate minerals with various Fe contents can help characterize the Fe concentration and oxidation state of the Earth’s interior. In the pyrolitic mantle, ringwoodite is thought to be the most abundant phase in the lower part of the mantle transition zone (Irifune and Ringwood, 1987). To understand the dynamics and structure of this deep mantle region, it is important to quantify the iron and water contents of possible transition zone materials. For this reason, some researchers have measured electrical conductivity of ringwoodite (Xu et al., 1998; Huang et al., 2005). Recently Yoshino et al. (2008a) measured both small polaron and proton conduction mechanisms of (Mg0.91 ,Fe0.09 )2 SiO4 wadsleyite and ringwoodite with various water contents. They suggested that a dry mantle model explains well the one-dimensional conductivitydepth profile obtained from the current electromagnetic field data beneath the Pacific Ocean (Kuvshinov et al., 2005). However, the profiles beneath the continents such as the Canadian Shield (Schultz et al., 1993; Neal et al., 2000) and Europe (Olsen, 1998; Tarits et al., 2004), yield lower conductivity values in the transition
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zone compared with estimates from the dry mantle model and a composition of Mg/(Mg + Fe) = 0.91. This pattern is especially evident in the lower part of the transition zone. If these Earth interior regions have a normal geotherm, chemical compositions beneath these continents could be different from the typical pyrolite composition. Thus, a better understanding of the effects of Fe content of ringwoodite on electrical conductivity is essential for constructing a reliable mineralogical model of the Earth’s mantle transition zone. In this study, we determined electrical conductivities of a series of ringwoodite samples with molar ratios of Fe/(Fe + Mg) = 0.09, 0.20 and 0.30, at a pressure of 20 GPa and temperatures ranging from 400 to 1900 K, in a Kawai-type multi-anvil apparatus. Using these experimental results, and in combination with geophysically reported, conductivity-depth profile data of the transition zone, we evaluated the composition of the lower part of the mantle transition zone. 2. Experimental methods To elucidate the effect of Fe contents on electrical conductivity of ringwoodite, we prepared starting materials with three different compositions of (Mg0.91 ,Fe0.09 )2 SiO4 , (Mg0.80 ,Fe0.20 )2 SiO4 and (Mg0.70 ,Fe0.30 )2 SiO4 . A ringwoodite sample with a Mg/(Mg + Fe) = 0.91 composition was synthesized from San Carlos olivine powder before the conductivity measurement. The powder was loaded in Mo capsule in a 6 mm edge length multi-anvil assembly and sintered at 20 GPa and 1800 K for 2.5 h to produce a pure ringwoodite with a grain size of several microns. The recovered sample was confirmed to be ringwoodite by X-ray diffraction, using a micro-beam X ray diffractometer in reflection geometry. The sample was then polished to remove the iron-poor part close
to the buffer material (Mo foil). A 1-mm core was removed from the center portion of the sintered ringwoodite aggregate sample with an ultrasonic drilling machine. For the conductivity measurement, the disk was placed in a capsule made of a single MgO crystal with inner diameter of 1 mm. Molybdenum electrodes with a diameter of 1 mm were placed in contact with the sample. The oxygen fugacity was controlled by the Mo/MoO2 buffer, which is close to the iron-wüstite buffer (Xu et al., 1998). Two sets of WRe3 –WRe25 thermocouples were mechanically connected to each Mo electrode in contact with the sample. The thermocouples were insulated from the LaCrO3 furnace by Al2 O3 and MgO insulators, and were taken out of the cell through the gaskets. The assemblage consisted of a Cr2 O3 -bearing MgO pressure medium, ZrO2 thermal insulator and a cylindrical LaCrO3 furnace. WC cubes with 3 mm truncation edge length were used as the inner anvils. The cell design for the conductivity measurements is shown in Fig. 1. In situ conductivity measurement was conducted using a Kawaitype multi-anvil apparatus. Electrical conductivity was measured by means of a pseudo 4-wire method with alternating current signal with an amplitude of 1 V and frequency of 0.1 Hz. Details for this procedure are described elsewhere (Fuji-ta et al., 2004; Yoshino et al., 2004). The samples were once heated to 1600 K and cooled to room temperature while measuring the conductivity. In subsequent heating, reversibility was confirmed. During each cycle, temperature was changed in 50–100 K steps, and electrical conductivity was measured at each temperature step. For conductivity measurements of ringwoodites with other chemical compositions, olivine powders with compositions of Fo80 and Fo70 were directly used without pre-synthesis. The olivine powder was filled in a MgO sleeve with an inner diameter of 1 mm and 1 mm thickness. The cell assembly for conductivity measure-
Fig. 1. Schematic cross-section of high-pressure cell assembly for conductivity measurements in Kawai-type multi-anvil press.
T. Yoshino, T. Katsura / Physics of the Earth and Planetary Interiors 174 (2009) 3–9
ments was the same as in the case of Fo91 . The samples were heated to 1500 K and annealed for 20 min to synthesize the ringwoodite aggregates. After annealing, conductivity measurements were taken during the first cooling and subsequent heating cycles. After annealing at 1500 K, electrical conductivities were reproducible. After the conductivity measurements were completed, the samples were slowly decompressed at room temperature for recovery. Phases in all the recovered samples were confirmed to be ringwoodite by X-ray diffraction. The samples were polished to observe the microstructures with a scanning electron microscope. At the contact between sample and electrode, no iron loss to the Mo electrode was observed. Sample geometry was largely changed due to compression by the hard Al2 O3 thermocouple insulators, which indented toward the Mo electrode when the olivine powder was used as a starting material. The distance between two Mo disks was measured to calculate the sample conductivity. 3. Experimental results Fig. 2 shows electrical conductivity of Fe-bearing ringwoodites as a function of reciprocal temperature. All samples examined in this study behaved as a semiconductor. The absolute conductivity () increased with increasing total Fe content, and the temperature dependence can be expressed according to the Arrhenian relation:
H
= 0 exp −
kT
(1)
where 0 is the pre-exponential factor, H is the activation enthalpy, k is the Boltzmann constant, and T is the absolute temperature. For each sample, the Arrhenius relations can be divided into high temperature and low temperature regimes. The pre-exponential term and activation enthalpy for each temperature regime are summarized in Table 1. The activation enthalpy of each sample is larger in the high than in the low temperature region. In the high temperature region, the activation enthalpy decreases with increasing total Fe content from 1.44 to 0.92 eV. It also decreases from 1.15 to 0.74 eV in the low temperature region. The temperature at which the activation enthalpy changes, slightly increases with decreasing the total Fe contents. Thus, the effect of Fe content is greater in the low temperature region. The pre-exponential terms determined at high temperatures are generally higher than those determined at
Fig. 2. Logarithm of electrical conductivity versus reciprocal temperature for ironbearing ringwoodites. Thick shaded line denotes the Arrhenius plot of ringwoodite polycrystals with 0.3 wt% H2 O (Xu et al., 1998).
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Table 1 Summary of runs. Run No.
Sample
T (K)
(S/m)
E (eV)
5K1035
Mg#70
400–1100 1150–1400
2.51(2) 3.30(3)
0.74(0) 0.92(0)
5K1022
Mg#80
550–1100 1200–1900
1.58(9) 3.45(4)
0.86(1) 1.26(1)
5K1047
SC Mg#91
650–850 900–1600
1.63(20) 3.31(1)
1.15(3) 1.44(0)
All experiments were conducted at 20 GPa.
low temperatures (Table 1). A previous study of the electrical conductivity of ringwoodites with a composition of Mg# = 90 (Xu et al., 1998) shows considerably higher conductivity values than the ringwoodites with a composition of Mg# = 91 of this study. This difference is due to the fact that their samples contained a significant amount of water (Huang et al., 2005). 4. Discussion 4.1. Charge transport mechanisms Electronic conduction in silicate minerals is thermally activated with energy barriers for formation and migration of charge carriers. Possible electron conduction mechanisms in ringwoodite include ionic conductivity with magnesium vacancies, large polaron and small polaron electronic conductivities. In the present study, the change in slope in the Arrhenius plot may be attributed to the conduction mechanism with temperature. Ionic conduction may be dominant as a conduction mechanism for mantle minerals in high temperature ranges (e.g., >1600 K for olivine; Schock et al., 1989). However, for ringwoodite, the temperature at which activation enthalpy changes, is much lower (900–1200 K). In addition, activation enthalpy in high temperature regimes is much lower than the >2 eV of ionic conduction obtained from other silicate minerals. A large polaron is another candidate of conduction mechanism at high temperatures. In fact, large polarons in magnesiowüstite have been proposed as a dominant conduction mechanism at high temperatures (>1000 K) (Li and Jeanloz, 1994; Dobson et al., 1997). In this case, the charge carriers are electron holes produced in the O 2p valence band and charge transfer occurs between O2− and Fe3+ . The mobility of large polaron charge carriers yields the temperature dependence of the pre-exponential factor ( 0 ) of T1/4 (Dobson and Brodholt, 2000). However, the present data show linearity at high temperature in an Arrhenius plot, which is not consistent with large polaron conduction. The present study demonstrates that electrical conductivities of ringwoodite exhibit strong Fe-content dependence. Similar studies concerning electrical conductivity as a function of Mg/(Mg + Fe) have been carried out for olivine (Hinze et al., 1981; Omura et al., 1989; Hirsch et al., 1993), pyroxenes (Seifert et al., 1982) and garnet (Romano et al., 2006). For all case, the absolute conductivity values increase with increasing total Fe contents under the same oxygen buffer. The Fe-content dependence of conductivities suggests small polaron conduction in the iron-bearing ringwoodites. If small polarons are the dominant conduction mechanism for both low and high temperature ranges, the change in activation enthalpy with temperature would be due to a change from metastable, quenchedin defect formation to a defect population, which is re-equilibrating with temperature. For small polaron conduction, charge transfers occur by electron–hole hopping between ferric and ferrous ions: • x FeMe + h˙ = FeMe
(2)
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The pre-exponential factor ( 0 ) in Eq. (1) should be defined as a function of iron content. It is known from the Nernst–Einstein equation that electrical conductivity depends on the number (N) of electric charge carriers per unit volume; = Nze
(3)
where z is the charge number, e is the electron charge, and is the mobility. The number of the charge carrier (N) is proportional to the ferric iron concentration in ringwoodite when Fe3+ / Fe is less than 0.5. Assuming the constant Fe3+ / Fe is independent of oxygen fugacity as a function of temperature, the electrical conductivity of iron-bearing ringwoodite would be proportional to the mole fraction of Fe in the Mg site (XFe );
H
= 1 XFe exp −
(4)
kT
where 1 is the constant. At high temperatures, formation of ferric iron is thermally activated. The activation enthalpy of ferric iron formation (charge balanced defect group) must be high enough to produce significant ferric iron concentration. However, the present results show that the difference in activation enthalpy between the low temperature and the high temperature regimes is too small (<0.3 eV). Therefore, the ferric iron concentration seems to be primarily controlled by the oxygen fugacity (fO2 ). For a single iron content (XFe ), 1 in Eq. (4) can be expressed as a function of oxygen fugacity;
H
1 = 2 fOn exp − 2
kT
(5)
where 2 is the constant. Fe3+ / Fe may be constant at low temperatures where the ferric iron content in ringwoodite no longer equilibrates with the oxygen buffer (Mo/MoO2 ) due to sluggish kinetics. In contrast, the equilibrium of ferric iron content may increase with temperature along the oxygen buffer line in the high temperature regime. As temperature increases, the increase of Fe3+ / Fe may lead to a relative increase in conductivity compared with the constant oxygen fugacity conditions. Fig. 3 shows activation enthalpies for electrical conductivity in iron-bearing ringwoodite plotted as a function of composi-
Fig. 4. Iron content versus activation enthalpy for small polaron conduction of ringwoodite. Solid and dashed lines indicate the best fitting line based on Eqs. (4) and (5), respectively. The correlation between iron content and activation enthalpy can be approximated well by a power of 1/3.
tion [Mg# = Mg/(Mg + Fe) × 100] in both high- and low-temperature regions. As previously described, activation enthalpies are systematically higher at high temperatures. The activation enthalpy, with increasing Mg# from 70 to 91, decreases with iron content from 1.44 to 0.92 eV and from 1.15 to 0.74 eV in the high and low temperature regions, respectively. The variation of activation enthalpies of ringwoodite at high temperatures is quite similar to those of olivine along the forsterite-fayalite join (Omura et al., 1989), of pyroxene along the ferrosilite–enstatite join (Seifert et al., 1982) and of garnet along the almandine-pyrope join at 10 GPa (Romano et al., 2006). On the other hand, electrical conductivity measurements of olivine with Mg# ranging from 66.5 to 91.35 at atmospheric pressure showed no clear dependence of activation enthalpy on iron content (Hirsch et al., 1993). However, the latter narrow temperature range in this previous study (1423–1573 K) might have prevented a reliable determination of activation enthalpies. Assuming the small polaron is the dominant mechanism for iron-bearing ringwoodite, a decrease in activation enthalpy with total Fe concentration is considered to be associated with a decrease in the average Fe3+ –Fe2+ distance. If the average Fe3+ –Fe2+ distance is isotropic, the activation enthalpy is expected to be proportional to the cube root of the total Fe content. This relationship can be approximated by an equation for the n-type semiconductor (Debye and Conwell, 1954), H = H0 − ˛(XFe )1/3
(6)
where XFe is the mole fraction of iron in the Mg site, H is the activation enthalpy at a certain value of XFe , H0 is the activation enthalpy observed at very low Fe concentrations, and ˛ is a constant accounting for geometrical factors. As shown in Fig. 4, the relationship between activation enthalpy and XFe is well fitted by Eq. (6), using the fitting parameters given in Table 2. The parameters fitted by Eq. (6) are as follows: H0 = 2.48 ± 0.42 eV and 1.98 ± 0.11 eV,
Fig. 3. Activation enthalpies for electrical conductivity in iron-bearing ringwoodite plotted as a function of composition Mg#. Open and closed circles represent activation enthalpy for low and high temperature regimes, respectively. Open and closed squares denote activation enthalpies for olivine at 4 GPa taken from Omura et al. (1989) and garnet at 10 GPa taken from Romano et al. (2006), respectively. Dashed line indicates activation enthalpy for pyroxenes taken from Seifert et al. (1982).
Table 2 Parameters fitted by Eq. (7).
High T Low T
A
H0 (eV)
˛
10042(4211) 467(150)
2.08(6) 2.14(5)
1.55(9) 2.14(7)
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Fig. 5. Electrical conductivity of iron-bearing ringwoodite as a function of total iron content for several temperatures. Symbols denote the measured electrical conductivity shown in Fig. 2. Solid lines indicate electrical conductivity of various iron content calculated from Eq. (5) as a summation of electrical conductivities of low and high temperature regimes.
and ˛ = 2.25 ± 0.72 and 1.88 ± 0.19 for high and low temperatures, respectively. Taking into account the concentration dependence of the pre-exponential factor as defined by Eq. (3), the resultant electrical conductivities of iron-bearing ringwoodite can be expressed as follows:
= AXFe exp
H0 − ˛XFe 1/3 − kT
(7)
where A is a constant including the dependence of oxygen fugacity on mole fraction of ferric iron as defined by Eq. (5). The constant A in Eq. (7) for the lower temperature regime does not depend on oxygen fugacity, whereas that of the high temperature regime would reflect a change of oxygen fugacity as shown in Eq. (5). The pre-exponential factor ( 0 ) in Eq. (1) determined in the low temperature region clearly increases with the total Fe content as is expected from the above discussion. On the other hand, the pre-exponential factor determined in the high temperature region is almost constant and independent of the total Fe content (Table 1). The constant 0 would suggest that the ratio Fe3+ / Fe under the same oxygen buffer conditions decreases with increasing XFe . The present data for low and high temperature regimes are fitted separately by Eq. (7) with the fitting parameters given in Table 2, because the pre-exponential terms in Eq. (7) for two regimes may not be equivalent as discussed above. Variations of conductivity values and activation enthalpy as a function of the total Fe content determined from Eq. (7) can explain the present results (Figs. 4 and 5). 4.2. Implications for the conductivity-depth profiles in the transition zone To estimate the conductivity-depth profile as a function of Fe content in ringwoodite, we consider a depth range of the transition zone between 520 and 660 km. Effects of additional phases (especially majorite) on the bulk rock conductivity were ignored to simplify the model, although the electrical conductivity of majorite garnet with a composition of pyrolite minus olivine is slightly lower than that of ringwoodite (Mg#91; Yoshino et al., 2008b). The tem-
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Fig. 6. Electrical conductivity profiles beneath the Pacific Ocean between 200 and 800 km depth. The shaded areas surrounded by solid line represent geophysically observed conductivity profiles beneath the Pacific Ocean (Kuvshinov et al., 2005). Thick light and dark gray lines denote the conductivity-depth profiles beneath the Canadian shield (Neal et al., 2000) and Europe (Tarits et al., 2004), respectively. Thick solid line represents the electrical conductivity of olivine, wadsleyite and ringwoodite with XFe = 0.91 (Yoshino et al., 2008a). Thin solid lines indicate the electrical conductivity of ringwoodite with various Fe contents. The thick dashed line indicates the electrical conductivity value of ringwoodite with Mg# = 90 for a decrease of 300 K in temperature compared to the normal geotherm.
perature profile in the transition zone was assumed to be mantle adiabat (Katsura et al., 2004). The oxygen fugacity of the mantle transition zone was assumed to be the iron–wüstite buffer (McCammon, 2005). Because the Mo–MoO2 buffer is thought to be close to the iron–wüstite buffer (Xu et al., 1998), the present conductivity data, based on Eq. (5), were directly applied without correction for oxygen fugacity. Because the expected temperature range of the mantle transition zone falls in the high temperature regime, conductivity values were calculated using the fitting parameters determined from this regime. Pressure dependence on electrical conductivity of ringwoodite was ignored. Fig. 6 illustrates the conductivity-depth profile in the transition zone, based on the present laboratory data. Iron content in ringwoodite considerably affects the conductivity structure of the transition zone. As the total Fe content in ringwoodite increases from XFe = 0.05 to 0.3, the conductivity value significantly increases by two orders of magnitude. When Fe content is very low (XFe < 0.1), the electrical conductivity of ringwoodite becomes more sensitive to the Fe content, because activation enthalpy is proportional to the cubic root of the Fe content. To estimate Fe content of ringwoodite in the transition zone, some geophysical observations are compared with our model. Utada et al. (2003) proposed a semi-global reference model for electrical conductivity based on electromagnetic response functions estimated from electric field variations and geomagnetic field variations beneath the Pacific Ocean, which covers a quarter of the Earth’s transition zone. Later, Kuvshinov et al. (2005) refined the reference model based on thorough consideration of the effect of the ocean on C-responses, and suggested that the mantle transition zone is more resistive than the previous model (Utada et al., 2003). The latter model is firstly considered as a representative conductivity-depth profile of the Earth’s upper mantle. Yoshino et al. (2008a) suggested that electrical conductivities of wadsleyite and ringwoodite with a composition of Mg# = 91 are consistent with the conductivity jumps at 410 and 660 km discontinuities. The present study also demonstrates that the conductivity values in the stability field of ringwoodite estimated from Mg# = 90 are quite consistent with those predicted from the electromagnetic observations. As shown in Fig. 6, high-pressure polymorphs of olivine in
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the mantle transition zone beneath the Pacific Ocean would have similar Fe content around Mg# = 90. The mantle transition zones beneath continent or stable craton (Europe and Canadian shield) have lower conductivity values than those beneath the Pacific Ocean in the stability field of ringwoodite (Schultz et al., 1993; Olsen, 1998; Neal et al., 2000; Tarits et al., 2004). There are two factors leading to lower conductivity values at these regions. One interpretation is that the temperature at the depth of the ringwoodite stability field beneath the continent is considerably lower than that beneath the Pacific Ocean. In this part of the mantle, a temperature of 300 K less than the normal geotherm explains the magnetotelluric conductivity profile of these regions. Tarits et al. (2004) proposed the presence of a dry and cool descending slab beneath southern Europe as the cause of the low conductivity values. In addition, Semenov and Jozwiak (2006) showed that the mantle conductance tends to be lower in southern Europe including the Mediterranean subduction zone, than the northern part near the Fennoscandia craton. This suggests the possibility that the subduction-driven low temperature is more effective in lowering the transition zone conductivity than chemical heterogeneity related to the surface geology. Another interpretation is that Fe content in ringwoodite at the depths of the ringwoodite stability field beneath these continents is significantly lower than that beneath the Pacific Ocean. If these regions have a normal geotherm, the Mg number of ringwoodite can be estimated to be around 95. This abnormally high Mg number requires the presence of a strongly depleted mantle, such as the continental tectosphere (Jordan, 1988). There is no seismic evidence to extend a continental root to the lower part of the mantle transition zone. In the stability field of wadsleyite, the laboratory-based conductivity values are quite similar to the geophysical observations. Conversely, the observed conductivity values in the stability field of ringwoodite are definitely lower compared with those obtained from our model. This disagreement may be caused by the presence of a stagnant slab just above the 660 km discontinuity. Thus, the harzbergite or depleted mantle composing the main part of the subducting slab would have a higher Mg number compared with the surrounding mantle. However, the predicted Mg number 91 of ringwoodite in hertzbergite is considerably lower than Mg# = 95 (Irifune and Ringwood, 1987). The global seismic tomography data beneath the regions just above the 660 km discontinuity generally show a high velocity anomaly (∼1% higher than the reference model) (e.g., Fukao et al., 2001; Zhao, 2004), suggesting higher Mg content in constituent minerals or lower temperatures compared with the surrounding mantle. Assuming the Mg# = 95 of ringwoodite, as estimated by electrical conductivity, P wave velocity increases 1.5% against the surrounding mantle (Higo et al., 2006). On the other hand, a decrease in temperature of 300 K compared with the normal geotherm leads to a 1% increment of the P wave velocity (Irifune et al., 2008). The high velocity anomaly at the bottom of the transition zone beneath the stable craton is quite consistent with a decrease of temperature of 300 K, as well as the low conductivity anomaly at that depth. Consequently, relatively low conductivity in these regions can be explained by the existence of a cooler region compared with the surrounding mantle, rather than the presence of iron-poor ringwoodite. If ringwoodite in the stagnant slab is slightly iron-poor, then a difference in temperature between the slab and the surrounding mantle would be smaller. Acknowledgments We are grateful to E. Ito, D. Yamazaki and H. Utada for discussion. This work was supported by a Grant-in-Aid for Scientific Research from the Japan Society for the Promotion of Science, No. 18740280
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Yoshino, T., Nishi, M., Matsuzaki, T., Yamazaki, D., Katsura, T., 2008b. Electrical conductivity of majorite garnet and its implications for electrical structure in the mantle transition zone. Phys. Earth Planet. Int., doi:10.1016/j.pepi. 2008.04.009. Zhao, D., 2004. Global tomographic images of mantle plumes and subducting slabs: insight into deep earth dynamics. Phys. Earth Planet. Int. 146, 3–34.
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Effect of iron content on electrical conductivity of ringwoodite, with implications for electrical structure in the transition zone Takashi Yoshino ∗ , Tomoo Katsura Institute for Study of the Earth’s Interior, Okayama University, Misasa, Tottori 682-0193, Japan
a r t i c l e
i n f o
Article history: Received 2 October 2007 Received in revised form 26 August 2008 Accepted 21 September 2008 Keywords: Electrical conductivity Iron content Ringwoodite Transition zone Upper mantle
a b s t r a c t Electrical conductivity of ringwoodite with various iron contents [Fe/(Fe + Mg) = 0.09, 0.2 and 0.3] was measured at pressure (20 GPa) and temperature (up to 1900 K) conditions of the lower part of the mantle transition zone in a Kawai-type multi-anvil apparatus. The conductivity increased with increasing total iron content. All electrical conductivity data were fitted to the formula of electrical conductivity = 0 XFe exp(−H/kT), where 0 is the pre-exponential term, XFe is the mole fraction of iron content in the Mg site, H is the activation enthalpy, k is the Boltzmann constant and T is absolute temperature. The activation enthalpy becomes higher at a certain temperature. At high temperatures, the activation enthalpy decreased from 1.44 to 0.92 eV with increasing total Fe content. At low temperatures less than 1000 K, the activation enthalpy also decreases from 1.15 to 0.74 eV with total Fe content. Dependence of the activation enthalpy on Fe content suggests that the dominant mechanism of charge transport is Fe2+ –Fe3+ hopping (small polaron). Recent electrical conductivity-depth profiles of the transition zone beneath the Pacific Ocean obtained from the electromagnetic induction study shows that the conductivity values between 520 and 660 km depths may be explained by ringwoodite with Fe/(Fe + Mg) = 0.10. On the other hand, assuming a normal geotherm, conductivity values beneath the continent or stable craton are considerably lower than those of ringwoodite with Fe/(Fe + Mg) = 0.10. Taking into consideration results from the global seismic tomographic studies, relatively low conductivity in these regions can be explained by the existence of a cooler region compared with the surrounding mantle, rather than the presence of iron-poor ringwoodite, or a combination of both. © 2008 Elsevier B.V. All rights reserved.
1. Introduction Electrical conductivity can be used to quantify temperature, water content, mineralogy, chemical composition and oxygen fugacity of the Earth’s deep mantle. In situ laboratory measurements of electrical conductivity of the major mantle phases under high-pressure and high-temperature conditions in conjunction with conductivity-depth profiles of the mantle derived from electromagnetic data provide some constraints on the temperature and chemistry in the Earth’s interior (e.g., Katsura et al., 1998; Xu et al., 1998; Yoshino et al., 2006, 2008a). For the iron-bearing silicate minerals under high-temperature conditions corresponding to the Earth’s deep interior, a small polaron conduction (hopping of electron holes between Fe2+ and Fe3+ ) has been considered as a dominant conduction mechanism (e.g., Schock et al., 1989; Hirsch et al., 1993; Katsura et al., 1998; Xu et al., 1998; Yoshino et al., 2006, 2008a). For this mechanism, the
∗ Corresponding author. Tel.: +81 858 34 3737; fax: +81 858 34 3450. E-mail address: [email protected] (T. Yoshino). 0031-9201/$ – see front matter © 2008 Elsevier B.V. All rights reserved. doi:10.1016/j.pepi.2008.09.015
conductivity of iron-bearing silicate minerals should be strongly controlled by the total Fe content and Fe3+ / Fe (i.e., oxidation condition). Thus, knowledge of electrical properties of silicate minerals with various Fe contents can help characterize the Fe concentration and oxidation state of the Earth’s interior. In the pyrolitic mantle, ringwoodite is thought to be the most abundant phase in the lower part of the mantle transition zone (Irifune and Ringwood, 1987). To understand the dynamics and structure of this deep mantle region, it is important to quantify the iron and water contents of possible transition zone materials. For this reason, some researchers have measured electrical conductivity of ringwoodite (Xu et al., 1998; Huang et al., 2005). Recently Yoshino et al. (2008a) measured both small polaron and proton conduction mechanisms of (Mg0.91 ,Fe0.09 )2 SiO4 wadsleyite and ringwoodite with various water contents. They suggested that a dry mantle model explains well the one-dimensional conductivitydepth profile obtained from the current electromagnetic field data beneath the Pacific Ocean (Kuvshinov et al., 2005). However, the profiles beneath the continents such as the Canadian Shield (Schultz et al., 1993; Neal et al., 2000) and Europe (Olsen, 1998; Tarits et al., 2004), yield lower conductivity values in the transition
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zone compared with estimates from the dry mantle model and a composition of Mg/(Mg + Fe) = 0.91. This pattern is especially evident in the lower part of the transition zone. If these Earth interior regions have a normal geotherm, chemical compositions beneath these continents could be different from the typical pyrolite composition. Thus, a better understanding of the effects of Fe content of ringwoodite on electrical conductivity is essential for constructing a reliable mineralogical model of the Earth’s mantle transition zone. In this study, we determined electrical conductivities of a series of ringwoodite samples with molar ratios of Fe/(Fe + Mg) = 0.09, 0.20 and 0.30, at a pressure of 20 GPa and temperatures ranging from 400 to 1900 K, in a Kawai-type multi-anvil apparatus. Using these experimental results, and in combination with geophysically reported, conductivity-depth profile data of the transition zone, we evaluated the composition of the lower part of the mantle transition zone. 2. Experimental methods To elucidate the effect of Fe contents on electrical conductivity of ringwoodite, we prepared starting materials with three different compositions of (Mg0.91 ,Fe0.09 )2 SiO4 , (Mg0.80 ,Fe0.20 )2 SiO4 and (Mg0.70 ,Fe0.30 )2 SiO4 . A ringwoodite sample with a Mg/(Mg + Fe) = 0.91 composition was synthesized from San Carlos olivine powder before the conductivity measurement. The powder was loaded in Mo capsule in a 6 mm edge length multi-anvil assembly and sintered at 20 GPa and 1800 K for 2.5 h to produce a pure ringwoodite with a grain size of several microns. The recovered sample was confirmed to be ringwoodite by X-ray diffraction, using a micro-beam X ray diffractometer in reflection geometry. The sample was then polished to remove the iron-poor part close
to the buffer material (Mo foil). A 1-mm core was removed from the center portion of the sintered ringwoodite aggregate sample with an ultrasonic drilling machine. For the conductivity measurement, the disk was placed in a capsule made of a single MgO crystal with inner diameter of 1 mm. Molybdenum electrodes with a diameter of 1 mm were placed in contact with the sample. The oxygen fugacity was controlled by the Mo/MoO2 buffer, which is close to the iron-wüstite buffer (Xu et al., 1998). Two sets of WRe3 –WRe25 thermocouples were mechanically connected to each Mo electrode in contact with the sample. The thermocouples were insulated from the LaCrO3 furnace by Al2 O3 and MgO insulators, and were taken out of the cell through the gaskets. The assemblage consisted of a Cr2 O3 -bearing MgO pressure medium, ZrO2 thermal insulator and a cylindrical LaCrO3 furnace. WC cubes with 3 mm truncation edge length were used as the inner anvils. The cell design for the conductivity measurements is shown in Fig. 1. In situ conductivity measurement was conducted using a Kawaitype multi-anvil apparatus. Electrical conductivity was measured by means of a pseudo 4-wire method with alternating current signal with an amplitude of 1 V and frequency of 0.1 Hz. Details for this procedure are described elsewhere (Fuji-ta et al., 2004; Yoshino et al., 2004). The samples were once heated to 1600 K and cooled to room temperature while measuring the conductivity. In subsequent heating, reversibility was confirmed. During each cycle, temperature was changed in 50–100 K steps, and electrical conductivity was measured at each temperature step. For conductivity measurements of ringwoodites with other chemical compositions, olivine powders with compositions of Fo80 and Fo70 were directly used without pre-synthesis. The olivine powder was filled in a MgO sleeve with an inner diameter of 1 mm and 1 mm thickness. The cell assembly for conductivity measure-
Fig. 1. Schematic cross-section of high-pressure cell assembly for conductivity measurements in Kawai-type multi-anvil press.
T. Yoshino, T. Katsura / Physics of the Earth and Planetary Interiors 174 (2009) 3–9
ments was the same as in the case of Fo91 . The samples were heated to 1500 K and annealed for 20 min to synthesize the ringwoodite aggregates. After annealing, conductivity measurements were taken during the first cooling and subsequent heating cycles. After annealing at 1500 K, electrical conductivities were reproducible. After the conductivity measurements were completed, the samples were slowly decompressed at room temperature for recovery. Phases in all the recovered samples were confirmed to be ringwoodite by X-ray diffraction. The samples were polished to observe the microstructures with a scanning electron microscope. At the contact between sample and electrode, no iron loss to the Mo electrode was observed. Sample geometry was largely changed due to compression by the hard Al2 O3 thermocouple insulators, which indented toward the Mo electrode when the olivine powder was used as a starting material. The distance between two Mo disks was measured to calculate the sample conductivity. 3. Experimental results Fig. 2 shows electrical conductivity of Fe-bearing ringwoodites as a function of reciprocal temperature. All samples examined in this study behaved as a semiconductor. The absolute conductivity () increased with increasing total Fe content, and the temperature dependence can be expressed according to the Arrhenian relation:
H
= 0 exp −
kT
(1)
where 0 is the pre-exponential factor, H is the activation enthalpy, k is the Boltzmann constant, and T is the absolute temperature. For each sample, the Arrhenius relations can be divided into high temperature and low temperature regimes. The pre-exponential term and activation enthalpy for each temperature regime are summarized in Table 1. The activation enthalpy of each sample is larger in the high than in the low temperature region. In the high temperature region, the activation enthalpy decreases with increasing total Fe content from 1.44 to 0.92 eV. It also decreases from 1.15 to 0.74 eV in the low temperature region. The temperature at which the activation enthalpy changes, slightly increases with decreasing the total Fe contents. Thus, the effect of Fe content is greater in the low temperature region. The pre-exponential terms determined at high temperatures are generally higher than those determined at
Fig. 2. Logarithm of electrical conductivity versus reciprocal temperature for ironbearing ringwoodites. Thick shaded line denotes the Arrhenius plot of ringwoodite polycrystals with 0.3 wt% H2 O (Xu et al., 1998).
5
Table 1 Summary of runs. Run No.
Sample
T (K)
(S/m)
E (eV)
5K1035
Mg#70
400–1100 1150–1400
2.51(2) 3.30(3)
0.74(0) 0.92(0)
5K1022
Mg#80
550–1100 1200–1900
1.58(9) 3.45(4)
0.86(1) 1.26(1)
5K1047
SC Mg#91
650–850 900–1600
1.63(20) 3.31(1)
1.15(3) 1.44(0)
All experiments were conducted at 20 GPa.
low temperatures (Table 1). A previous study of the electrical conductivity of ringwoodites with a composition of Mg# = 90 (Xu et al., 1998) shows considerably higher conductivity values than the ringwoodites with a composition of Mg# = 91 of this study. This difference is due to the fact that their samples contained a significant amount of water (Huang et al., 2005). 4. Discussion 4.1. Charge transport mechanisms Electronic conduction in silicate minerals is thermally activated with energy barriers for formation and migration of charge carriers. Possible electron conduction mechanisms in ringwoodite include ionic conductivity with magnesium vacancies, large polaron and small polaron electronic conductivities. In the present study, the change in slope in the Arrhenius plot may be attributed to the conduction mechanism with temperature. Ionic conduction may be dominant as a conduction mechanism for mantle minerals in high temperature ranges (e.g., >1600 K for olivine; Schock et al., 1989). However, for ringwoodite, the temperature at which activation enthalpy changes, is much lower (900–1200 K). In addition, activation enthalpy in high temperature regimes is much lower than the >2 eV of ionic conduction obtained from other silicate minerals. A large polaron is another candidate of conduction mechanism at high temperatures. In fact, large polarons in magnesiowüstite have been proposed as a dominant conduction mechanism at high temperatures (>1000 K) (Li and Jeanloz, 1994; Dobson et al., 1997). In this case, the charge carriers are electron holes produced in the O 2p valence band and charge transfer occurs between O2− and Fe3+ . The mobility of large polaron charge carriers yields the temperature dependence of the pre-exponential factor ( 0 ) of T1/4 (Dobson and Brodholt, 2000). However, the present data show linearity at high temperature in an Arrhenius plot, which is not consistent with large polaron conduction. The present study demonstrates that electrical conductivities of ringwoodite exhibit strong Fe-content dependence. Similar studies concerning electrical conductivity as a function of Mg/(Mg + Fe) have been carried out for olivine (Hinze et al., 1981; Omura et al., 1989; Hirsch et al., 1993), pyroxenes (Seifert et al., 1982) and garnet (Romano et al., 2006). For all case, the absolute conductivity values increase with increasing total Fe contents under the same oxygen buffer. The Fe-content dependence of conductivities suggests small polaron conduction in the iron-bearing ringwoodites. If small polarons are the dominant conduction mechanism for both low and high temperature ranges, the change in activation enthalpy with temperature would be due to a change from metastable, quenchedin defect formation to a defect population, which is re-equilibrating with temperature. For small polaron conduction, charge transfers occur by electron–hole hopping between ferric and ferrous ions: • x FeMe + h˙ = FeMe
(2)
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The pre-exponential factor ( 0 ) in Eq. (1) should be defined as a function of iron content. It is known from the Nernst–Einstein equation that electrical conductivity depends on the number (N) of electric charge carriers per unit volume; = Nze
(3)
where z is the charge number, e is the electron charge, and is the mobility. The number of the charge carrier (N) is proportional to the ferric iron concentration in ringwoodite when Fe3+ / Fe is less than 0.5. Assuming the constant Fe3+ / Fe is independent of oxygen fugacity as a function of temperature, the electrical conductivity of iron-bearing ringwoodite would be proportional to the mole fraction of Fe in the Mg site (XFe );
H
= 1 XFe exp −
(4)
kT
where 1 is the constant. At high temperatures, formation of ferric iron is thermally activated. The activation enthalpy of ferric iron formation (charge balanced defect group) must be high enough to produce significant ferric iron concentration. However, the present results show that the difference in activation enthalpy between the low temperature and the high temperature regimes is too small (<0.3 eV). Therefore, the ferric iron concentration seems to be primarily controlled by the oxygen fugacity (fO2 ). For a single iron content (XFe ), 1 in Eq. (4) can be expressed as a function of oxygen fugacity;
H
1 = 2 fOn exp − 2
kT
(5)
where 2 is the constant. Fe3+ / Fe may be constant at low temperatures where the ferric iron content in ringwoodite no longer equilibrates with the oxygen buffer (Mo/MoO2 ) due to sluggish kinetics. In contrast, the equilibrium of ferric iron content may increase with temperature along the oxygen buffer line in the high temperature regime. As temperature increases, the increase of Fe3+ / Fe may lead to a relative increase in conductivity compared with the constant oxygen fugacity conditions. Fig. 3 shows activation enthalpies for electrical conductivity in iron-bearing ringwoodite plotted as a function of composi-
Fig. 4. Iron content versus activation enthalpy for small polaron conduction of ringwoodite. Solid and dashed lines indicate the best fitting line based on Eqs. (4) and (5), respectively. The correlation between iron content and activation enthalpy can be approximated well by a power of 1/3.
tion [Mg# = Mg/(Mg + Fe) × 100] in both high- and low-temperature regions. As previously described, activation enthalpies are systematically higher at high temperatures. The activation enthalpy, with increasing Mg# from 70 to 91, decreases with iron content from 1.44 to 0.92 eV and from 1.15 to 0.74 eV in the high and low temperature regions, respectively. The variation of activation enthalpies of ringwoodite at high temperatures is quite similar to those of olivine along the forsterite-fayalite join (Omura et al., 1989), of pyroxene along the ferrosilite–enstatite join (Seifert et al., 1982) and of garnet along the almandine-pyrope join at 10 GPa (Romano et al., 2006). On the other hand, electrical conductivity measurements of olivine with Mg# ranging from 66.5 to 91.35 at atmospheric pressure showed no clear dependence of activation enthalpy on iron content (Hirsch et al., 1993). However, the latter narrow temperature range in this previous study (1423–1573 K) might have prevented a reliable determination of activation enthalpies. Assuming the small polaron is the dominant mechanism for iron-bearing ringwoodite, a decrease in activation enthalpy with total Fe concentration is considered to be associated with a decrease in the average Fe3+ –Fe2+ distance. If the average Fe3+ –Fe2+ distance is isotropic, the activation enthalpy is expected to be proportional to the cube root of the total Fe content. This relationship can be approximated by an equation for the n-type semiconductor (Debye and Conwell, 1954), H = H0 − ˛(XFe )1/3
(6)
where XFe is the mole fraction of iron in the Mg site, H is the activation enthalpy at a certain value of XFe , H0 is the activation enthalpy observed at very low Fe concentrations, and ˛ is a constant accounting for geometrical factors. As shown in Fig. 4, the relationship between activation enthalpy and XFe is well fitted by Eq. (6), using the fitting parameters given in Table 2. The parameters fitted by Eq. (6) are as follows: H0 = 2.48 ± 0.42 eV and 1.98 ± 0.11 eV,
Fig. 3. Activation enthalpies for electrical conductivity in iron-bearing ringwoodite plotted as a function of composition Mg#. Open and closed circles represent activation enthalpy for low and high temperature regimes, respectively. Open and closed squares denote activation enthalpies for olivine at 4 GPa taken from Omura et al. (1989) and garnet at 10 GPa taken from Romano et al. (2006), respectively. Dashed line indicates activation enthalpy for pyroxenes taken from Seifert et al. (1982).
Table 2 Parameters fitted by Eq. (7).
High T Low T
A
H0 (eV)
˛
10042(4211) 467(150)
2.08(6) 2.14(5)
1.55(9) 2.14(7)
T. Yoshino, T. Katsura / Physics of the Earth and Planetary Interiors 174 (2009) 3–9
Fig. 5. Electrical conductivity of iron-bearing ringwoodite as a function of total iron content for several temperatures. Symbols denote the measured electrical conductivity shown in Fig. 2. Solid lines indicate electrical conductivity of various iron content calculated from Eq. (5) as a summation of electrical conductivities of low and high temperature regimes.
and ˛ = 2.25 ± 0.72 and 1.88 ± 0.19 for high and low temperatures, respectively. Taking into account the concentration dependence of the pre-exponential factor as defined by Eq. (3), the resultant electrical conductivities of iron-bearing ringwoodite can be expressed as follows:
= AXFe exp
H0 − ˛XFe 1/3 − kT
(7)
where A is a constant including the dependence of oxygen fugacity on mole fraction of ferric iron as defined by Eq. (5). The constant A in Eq. (7) for the lower temperature regime does not depend on oxygen fugacity, whereas that of the high temperature regime would reflect a change of oxygen fugacity as shown in Eq. (5). The pre-exponential factor ( 0 ) in Eq. (1) determined in the low temperature region clearly increases with the total Fe content as is expected from the above discussion. On the other hand, the pre-exponential factor determined in the high temperature region is almost constant and independent of the total Fe content (Table 1). The constant 0 would suggest that the ratio Fe3+ / Fe under the same oxygen buffer conditions decreases with increasing XFe . The present data for low and high temperature regimes are fitted separately by Eq. (7) with the fitting parameters given in Table 2, because the pre-exponential terms in Eq. (7) for two regimes may not be equivalent as discussed above. Variations of conductivity values and activation enthalpy as a function of the total Fe content determined from Eq. (7) can explain the present results (Figs. 4 and 5). 4.2. Implications for the conductivity-depth profiles in the transition zone To estimate the conductivity-depth profile as a function of Fe content in ringwoodite, we consider a depth range of the transition zone between 520 and 660 km. Effects of additional phases (especially majorite) on the bulk rock conductivity were ignored to simplify the model, although the electrical conductivity of majorite garnet with a composition of pyrolite minus olivine is slightly lower than that of ringwoodite (Mg#91; Yoshino et al., 2008b). The tem-
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Fig. 6. Electrical conductivity profiles beneath the Pacific Ocean between 200 and 800 km depth. The shaded areas surrounded by solid line represent geophysically observed conductivity profiles beneath the Pacific Ocean (Kuvshinov et al., 2005). Thick light and dark gray lines denote the conductivity-depth profiles beneath the Canadian shield (Neal et al., 2000) and Europe (Tarits et al., 2004), respectively. Thick solid line represents the electrical conductivity of olivine, wadsleyite and ringwoodite with XFe = 0.91 (Yoshino et al., 2008a). Thin solid lines indicate the electrical conductivity of ringwoodite with various Fe contents. The thick dashed line indicates the electrical conductivity value of ringwoodite with Mg# = 90 for a decrease of 300 K in temperature compared to the normal geotherm.
perature profile in the transition zone was assumed to be mantle adiabat (Katsura et al., 2004). The oxygen fugacity of the mantle transition zone was assumed to be the iron–wüstite buffer (McCammon, 2005). Because the Mo–MoO2 buffer is thought to be close to the iron–wüstite buffer (Xu et al., 1998), the present conductivity data, based on Eq. (5), were directly applied without correction for oxygen fugacity. Because the expected temperature range of the mantle transition zone falls in the high temperature regime, conductivity values were calculated using the fitting parameters determined from this regime. Pressure dependence on electrical conductivity of ringwoodite was ignored. Fig. 6 illustrates the conductivity-depth profile in the transition zone, based on the present laboratory data. Iron content in ringwoodite considerably affects the conductivity structure of the transition zone. As the total Fe content in ringwoodite increases from XFe = 0.05 to 0.3, the conductivity value significantly increases by two orders of magnitude. When Fe content is very low (XFe < 0.1), the electrical conductivity of ringwoodite becomes more sensitive to the Fe content, because activation enthalpy is proportional to the cubic root of the Fe content. To estimate Fe content of ringwoodite in the transition zone, some geophysical observations are compared with our model. Utada et al. (2003) proposed a semi-global reference model for electrical conductivity based on electromagnetic response functions estimated from electric field variations and geomagnetic field variations beneath the Pacific Ocean, which covers a quarter of the Earth’s transition zone. Later, Kuvshinov et al. (2005) refined the reference model based on thorough consideration of the effect of the ocean on C-responses, and suggested that the mantle transition zone is more resistive than the previous model (Utada et al., 2003). The latter model is firstly considered as a representative conductivity-depth profile of the Earth’s upper mantle. Yoshino et al. (2008a) suggested that electrical conductivities of wadsleyite and ringwoodite with a composition of Mg# = 91 are consistent with the conductivity jumps at 410 and 660 km discontinuities. The present study also demonstrates that the conductivity values in the stability field of ringwoodite estimated from Mg# = 90 are quite consistent with those predicted from the electromagnetic observations. As shown in Fig. 6, high-pressure polymorphs of olivine in
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the mantle transition zone beneath the Pacific Ocean would have similar Fe content around Mg# = 90. The mantle transition zones beneath continent or stable craton (Europe and Canadian shield) have lower conductivity values than those beneath the Pacific Ocean in the stability field of ringwoodite (Schultz et al., 1993; Olsen, 1998; Neal et al., 2000; Tarits et al., 2004). There are two factors leading to lower conductivity values at these regions. One interpretation is that the temperature at the depth of the ringwoodite stability field beneath the continent is considerably lower than that beneath the Pacific Ocean. In this part of the mantle, a temperature of 300 K less than the normal geotherm explains the magnetotelluric conductivity profile of these regions. Tarits et al. (2004) proposed the presence of a dry and cool descending slab beneath southern Europe as the cause of the low conductivity values. In addition, Semenov and Jozwiak (2006) showed that the mantle conductance tends to be lower in southern Europe including the Mediterranean subduction zone, than the northern part near the Fennoscandia craton. This suggests the possibility that the subduction-driven low temperature is more effective in lowering the transition zone conductivity than chemical heterogeneity related to the surface geology. Another interpretation is that Fe content in ringwoodite at the depths of the ringwoodite stability field beneath these continents is significantly lower than that beneath the Pacific Ocean. If these regions have a normal geotherm, the Mg number of ringwoodite can be estimated to be around 95. This abnormally high Mg number requires the presence of a strongly depleted mantle, such as the continental tectosphere (Jordan, 1988). There is no seismic evidence to extend a continental root to the lower part of the mantle transition zone. In the stability field of wadsleyite, the laboratory-based conductivity values are quite similar to the geophysical observations. Conversely, the observed conductivity values in the stability field of ringwoodite are definitely lower compared with those obtained from our model. This disagreement may be caused by the presence of a stagnant slab just above the 660 km discontinuity. Thus, the harzbergite or depleted mantle composing the main part of the subducting slab would have a higher Mg number compared with the surrounding mantle. However, the predicted Mg number 91 of ringwoodite in hertzbergite is considerably lower than Mg# = 95 (Irifune and Ringwood, 1987). The global seismic tomography data beneath the regions just above the 660 km discontinuity generally show a high velocity anomaly (∼1% higher than the reference model) (e.g., Fukao et al., 2001; Zhao, 2004), suggesting higher Mg content in constituent minerals or lower temperatures compared with the surrounding mantle. Assuming the Mg# = 95 of ringwoodite, as estimated by electrical conductivity, P wave velocity increases 1.5% against the surrounding mantle (Higo et al., 2006). On the other hand, a decrease in temperature of 300 K compared with the normal geotherm leads to a 1% increment of the P wave velocity (Irifune et al., 2008). The high velocity anomaly at the bottom of the transition zone beneath the stable craton is quite consistent with a decrease of temperature of 300 K, as well as the low conductivity anomaly at that depth. Consequently, relatively low conductivity in these regions can be explained by the existence of a cooler region compared with the surrounding mantle, rather than the presence of iron-poor ringwoodite. If ringwoodite in the stagnant slab is slightly iron-poor, then a difference in temperature between the slab and the surrounding mantle would be smaller. Acknowledgments We are grateful to E. Ito, D. Yamazaki and H. Utada for discussion. This work was supported by a Grant-in-Aid for Scientific Research from the Japan Society for the Promotion of Science, No. 18740280
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