Orientation-related electrical conductivity of hydrous olivine, clinopyroxene and plagioclase and implications for the structure of the lower continental crust and uppermost mantle

Earth and Planetary Science Letters 317–318 (2012) 241–250

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Orientation-related electrical conductivity of hydrous olivine, clinopyroxene and plagioclase and implications for the structure of the lower continental crust and uppermost mantle Xiaozhi Yang Bayerisches Geoinstitut, Universität Bayreuth, D-95440 Bayreuth, Germany

a r t i c l e

i n f o

Article history: Received 2 August 2011 Received in revised form 31 October 2011 Accepted 16 November 2011 Available online 27 December 2011 Editor: L. Stixrude Keywords: electrical anisotropy and conductivity orientation hydrous olivine, clinopyroxene and plagioclase lower continental crust uppermost mantle

a b s t r a c t Orientation-related electrical conductivity of H-annealed olivine, clinopyroxene and plagioclase single crystals have been measured at 10 kbar and 200–800 °C with an end-loaded piston cylinder apparatus and a Solartron-1260 Impedance/Gain Phase analyzer in the frequency range of 106–0.1 Hz. The complex spectra usually show an arc over the whole frequency range at low temperature, with much scatter of data-points at low frequencies for samples of large resistance, and an arc plus a short tail in the high and low frequency range, respectively, at high temperature. The arc is due to grain interior conduction, and the short tail is related to electrode effects. The conduction is dominated by H-related point defects. The results show negligible anisotropy in electrical conductivity for hydrous olivine and clinopyroxene but large anisotropy for hydrous plagioclase. Electrical anisotropy in the lower continental crust may be caused by the fabrics of the constitutive granulites due to lattice-preferred orientation (LPO) of hydrous plagioclase and banded pyroxene- and plagioclase-rich microstructure. Electrical anisotropy in the uppermost mantle cannot be produced by hydrous olivine (and pyroxenes in peridotites), but may be related to mantle macro-heterogeneity due to the local presence of Fe 3 +- and H2O-rich augites and other pyroxenites in the form of dykes/veins and/or remnants of recycled crustal materials in the form of stretched lenses. © 2011 Elsevier B.V. All rights reserved.

1. Introduction Electrical anisotropy has been reported concomitantly with electrical anomalies and synonymously with seismic anisotropy in many regions of the lower continental crust, ~ 20–40 km depth by average, and uppermost oceanic and continental mantle, ~ 40–200 km depth, and the resolved electrical conductivity can be up to ~ 0.01–0.1 S/m along some directions (Bahr and Duba, 2000; Evans et al., 2005; Hamilton et al., 2006; Jones, 1992; Leibecker et al., 2002; Mareschal et al., 1995; Simpson, 2001). In the lower continental crust, these zones are usually also characterized by seismic reflectivity (Mooney and Meissner, 1992); while in the uppermost mantle, these zones are usually also featured by low seismic velocity (Evans et al., 2005; Fischer et al., 2010). Variation of seismic anisotropy is usually on the scale of several percent, by contrast, electrical anisotropy in the lower crust and uppermost mantle can vary by a factor of ~ 2 to >10 (Bahr and Duba, 2000; Evans et al., 2005; Hamilton et al., 2006; Leibecker et al., 2002; Mareschal et al., 1995; Simpson, 2001). Although many models have been proposed to explain the geophysically measured high electrical conductivity, e.g., by the presence of inter-connected highly conductive secondary phases such as sulfides,

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graphite films, melts or fluids (Bahr and Duba, 2000; Jones, 1992; Mareschal et al., 1995), the origin of electrical anisotropy remains a long-standing and controversial issue (Bahr and Duba, 2000; Evans et al., 2005; Simpson, 2002). The lower continental crust and uppermost mantle consist mainly of nominally anhydrous olivine, pyroxenes and plagioclase, and these minerals generally contain some water as H-related point defects (OH/molecular H2O), e.g., from less than 100 to more than 1000 ppm H2O (by weight: the same unit is used throughout this paper) (Bell and Rossman, 1992; Ingrin and Skogby, 2000; Xia et al., 2006; Yang et al., 2008). The presence of H can enhance conduction (Karato, 1990), and because H diffusivity is highly anisotropic in many silicate minerals (Ingrin and Blanchard, 2006), some scholars argued that, based on theoretical calculations using H self-diffusion coefficients, electrical anisotropy in the uppermost mantle may be attributed to LPO of hydrous olivine (Evans et al., 2005; Simpson, 2002). However, recent experimental studies have shown that the mobility of H in silicate minerals is probably different between diffusion and conduction (Dai and Karato, 2009a,b,c; Huang et al., 2005; Poe et al., 2010; Romano et al., 2009; Wang et al., 2006; Yang and McCammon, 2011; Yang et al., 2011a,b; Yoshino et al., 2009). Thus, diffusion-based conductivity anisotropy of hydrous silicates may not be applied to natural systems, and laboratory conductivity measurements along different orientations are necessary.

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X. Yang / Earth and Planetary Science Letters 317–318 (2012) 241–250

Table 1 Chemical composition of the samples (wt.%). Mineral

SiO2

TiO2

Al2O3

FeO

MnO

MgO

CaO

Na2O

K2O

Total

Olivine Clinopyroxene Plagioclase

39.81 47.97 52.10

0.00 1.48 0.07

0.01 9.57 29.32

8.81 8.26 0.39

0.12 0.12 0.01

50.72 12.14 0.07

0.04 17.57 11.95

0.00 2.25 4.57

0.00 0.00 0.25

99.52 99.36 98.73

FeO: with all Fe assumed as FeO. Data are the average of multi-points measurements (6–20) by electron microprobe.

In this study, we have systematically measured the orientationrelated electrical conductivity of single crystal hydrous olivine, clinopyroxene and plagioclase under well-constrained temperature, pressure, water content and oxygen fugacity conditions with impedance spectroscopy. The results are of fundamental implications for a better knowledge of the electrical anisotropy and high conductivity in the lower continental crust and uppermost mantle, as well as their other physical properties. 2. Experimental methods 2.1. Sample characterization The starting materials were gem-quality olivine crystals from Dak Lak (Vietnam), an augite crystal from Nushan (China) and a plagioclase crystal from Aksu (China). The composition is homogeneous for each mineral (Table 1). The initial H2O content was b1 ppm (defined as “dry” in this paper) for the starting olivine, ~ 60 ppm for the starting clinopyroxene and ~ 50 ppm for the starting plagioclase (by Fourier-transform infrared (FTIR) spectroscopy). The initial Fe 3 +/Fetotal ratio was 19 ± 3% for the starting clinopyroxene (by Mössbauer spectroscopy: Appendix A). Olivine and clinopyroxene crystals were oriented using an X-ray precession camera, and (010) plane of the plagioclase crystal was determined by the developed (010) cleavage. Cylindrical cores with a diameter of ~2.0 mm were prepared with a diamond drilling machine: for orthorhombic olivine and monoclinic clinopyroxene, samples were drilled along directions /(100), /(010) and /(001), respectively; for triclinic plagioclase, samples were drilled along three mutually perpendicular directions, of which one is /(010) and the other two are, for simplicity, termed as /(100)′ and /(001)′, respectively (Table 2). Electrical properties of silicate minerals are sensitive to water content and oxygen fugacity, and in order to address the H-related effect of orientation on conductivity, hydrous samples with the same starting material should be better annealed under the same water- and oxygen-fugacity conditions (if the initial water contents are very low). In addition, in order to minimize uncertainty from sample geometry on conductivity calculation, recovered crystals should be free of any fractures and as large as possible. The hydrous olivine, clinopyroxene and

plagioclase were annealed in rapid-quench cold-seal vessels with water as pressure media. For each mineral, the cylindrical crystals, together with ~20–50% distilled water and a small piece of unoriented crystal which was used to measure the water content of the run products, were loaded into a Au-capsule (OD 5.0 mm, ID 4.8 mm and length ~2.0–3.0 cm) and the capsule was then welded. For olivine and clinopyroxene, a piece of Ni metal was also loaded into the capsule to buffer the oxygen fugacity. The hydration experiments were carried out at 2 kbar and 800 °C for 175–190 h. After each run, the presence of excess water in the recovered capsule was confirmed. For the run of olivine and clinopyroxene, both Ni and NiO were present in the recovered capsules. The Fe3 +/Fetotal ratio was 0 ±2% and 6 ± 4% for the annealed olivine and clinopyroxene, respectively. The very low Fe3 + content in the annealed olivine is consistent with previous work that the solubility of Fe3 + in mantle olivines is usually negligible (Canil et al., 1994). The variation of Fe3 +/Fetotal ratio from the starting to annealed clinopyroxene, ~19% to 6%, may be related to the great incorporation of H (see discussion below). The recovered crystals demonstrated no fractures or changes in geometry, and were cut and polished into disk-shapes (diameter ~2.0 mm and length ~1.6–2.0 mm) and ultrasonically cleaned for complex impedance spectroscopy. 2.2. FTIR spectroscopy Intensity of H-species in silicate minerals measured by FTIR spectroscopy depends on the orientation of the IR active dipole relative to the incident radiation, and thus, an accurate quantitative measurement of H2O content in anisotropic minerals requires analyses of oriented crystals using polarized light along the three principal optical directions. However, Libowitzky and Rossman (1996) showed that three polarized spectra along any three mutually perpendicular directions can give the same result. This method was used to determine the water content of the H-annealed and recovered samples prior to and after electrical conductivity runs, respectively. IR spectra were recorded with a Bruker IFS 120 FTIR spectrometer coupled with a Bruker IR microscope, using a tungsten light source, a Si-coated CaF2 beam-splitter and a MCT detector. The radiation was polarized using a wire-strip polarizer on a KRS-5 substrate. 100/200 scans were conducted for each spectrum at a 4 cm − 1

Table 2 Summary of samples and fitting parameters. Mineral Olivine

Initial ppm H2O 40

Clinopyroxene

680

Plagioclase

105

Orientation

L/S (m− 1)

Final ppm H2O

Log10 (AΗ (S/m))

ΔHH (kJ/mol)

/(100) /(010) /(001) /(100) /(010) /(001) /(100)′ /(010) /(001)′

602 571 607 698 741 759 661 639 693

35 40 35 655 690 675 95 110 105

0.93 ± 0.26 0.93 ± 0.52 0.38 ± 0.47 0.71 ± 0.07 0.70 ± 0.07 0.58 ± 0.08 2.85 ± 0.24 3.97 ± 0.22 2.80 ± 0.34

92 ± 4 92 ± 7 84 ± 6 65 ± 1 63 ± 1 62 ± 1 124 ± 4 130 ± 4 115 ± 5

H2O data were rounded to the nearest 5 wt.ppm. The reported final water content was measured in the region apart from the outmost rims of the recovered crystals where slight dehydration occurred for some samples (see text). L/S, the effective ratio of length to cross section area of the recovered samples. Eq. (2) was used to obtain the fitting parameters. Uncertainty is one standard deviation.

X. Yang / Earth and Planetary Science Letters 317–318 (2012) 241–250

resolution, with spot sizes of 30–100 μm in diameter. The polarized spectra were measured with the electric field vector (E) parallel to three mutually perpendicular directions (X, Y and Z), respectively. The water content was calculated from Cw ¼

ΔX þ ΔY þ ΔZ I

ð1Þ

where Cw is the concentration (wt. ppm H2O), ΔX, Y, Z is the thicknessnormalized integral absorbance (cm − 1) of H-species along X, Y and Z directions (from 3700 to 2500 cm − 1), respectively, and I is the integral specific absorption coefficient (ppm − 1 cm − 2) from Bell et al. (2003) for olivine, from Bell et al. (1995) for clinopyroxene and from Johnson and Rossman (2003) for plagioclase. The uncertainty of H2O content is ~ 5–15%, owing mainly to background subtraction. 2.3. Complex impedance spectroscopy Because the effect of pressure, given a variation of several GPa, is negligible on the electrical conductivity of silicate minerals under either dry or hydrous conditions (Dai and Karato, 2009b; Xu et al., 2000), the complex spectra were measured at 10 kbar with an endloaded piston cylinder apparatus and a Solartron 1260 Impedance/ Gain Phase analyzer, with a voltage of 500 mV and a frequency range of 10 6–0.1 Hz in the StandAlone mode. The assembly setup is shown in Fig. 1 (see also Appendix B). By this method, resistance below the

243

order of 10 9 Ω can be determined accurately. Because the hydrous samples were annealed at 800 °C, the impedance was measured at 200–800 °C to avoid serious dehydration. For each run, after the designated pressure was reached, the system was first pre-heated at 200–300 °C for up to ~2 h. To minimize H-loss especially at high temperatures, the temperature was changed at a rate of 100 °C/min, and the impedance was analyzed at various temperatures upon heating/ cooling (prior to each analysis, the temperature was maintained for 2–10 min after the pre-heating). After each run, the recovered sample was polished for subsequent optical, Raman and FTIR spectroscopic examinations. No other phases were found, and both Ni and NiO were still present. Distortions of sample geometry were negligible (Fig. 2), although the absolute diameter and length decreased slightly. The sample dimensions after the run, considering the effective contact between the sample and electrodes, were used to calculate the electrical conductivity (σ), σ = L / SR, where L and S are the sample length and cross section area, respectively, and R is the resistance. The resistance was obtained by fitting the high frequency arc of the complex spectra (see below). During each measurement, the variation of temperature was within ± 1 °C, however, the total uncertainty, including that arising from thermal gradients along the length of the capsule and from the thermocouple itself, is b20 °C. The total uncertainty of conductivity, including sample dimensions and impedance arc fitting, is b20% (mostly b10%). 3. Results Polarized FTIR spectra of the annealed samples are shown in Fig. 3, and all the spectra show OH-related bands from 3700 to 2500 cm − 1 (Bell et al., 2004; Johnson and Rossman, 2004; Skogby et al., 1990). The H2O content is ~40,680 and 105 ppm for olivine, clinopyroxene and plagioclase, respectively (Table 2). The low water content of olivine agrees with the low solubility of H in natural olivines at low pressures and temperatures (Bai and Kohlstedt, 1993), and the water content of plagioclase is comparable to that reported for a natural orthoclase crystal H-annealed at similar conditions (Wang et al., 1997). In contrast, the water content of clinopyroxene is >10 times higher than that of a natural Fe-poor diopside at similar conditions (Bromiley et al., 2004). The decrease of Fe 3 +/Fetotal in the clinopyroxene, e.g., from ~ 19% in the starting material to ~6% in the annealed sample which is possible in hydration annealing even the oxygen fugacity of the system is buffered (Bromiley et al., 2004), indicates the dissolution of H by Ni + H2O = NiO + H2 and 2Fe 3 + + 2O 2 − + H2 = 2Fe 2 + + 2OH −.

Fig. 1. Sample assembly for electrical conductivity measurement. A Pt-capsule, with one end welded and the other one sealed by boron nitride (BN), was used, and the sample was enclosed by a tube-shape BN. Pt/Au disks (and Pt-capsule) were used as electrodes, and Ni–NiO was used as oxygen buffer. The use of Pt-capsule and BN produces a relatively sealed chamber, prevents large variation in water content by diffusion loss/gain and maintains the sample geometry, because BN is of low strength, high insulation and low permeability for H-species. A piece of Ni foil connected to ground was placed between the heater and the sample to minimize the electric disturbance from the heater, the temperature gradient in the sample chamber and the leakage current through the system, and a S-type thermocouple was used to measure the temperature. The component parts, such as the capsule, BN and Al2O3, were heated at 1000 °C for 3–6 h prior to sample assembly. After the assembly was completed, it was heated at ~ 150–200 °C in vacuum overnight before conductivity measurements started. No glue/cement was used to immobilize the experimental parts during assembly to avoid the release of volatiles upon heating.

Fig. 2. Cross-section of the recovered capsule from an electrical conductivity measurement on olivine /(001). Note that the fractures were produced mainly during the final decompression and post-experimental sample polishing.

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X. Yang / Earth and Planetary Science Letters 317–318 (2012) 241–250

2

a

b

c

Absorbance

24

8

18

1

6

12

0 4000

3800

3600

3400

3200

3000

4000

Wavenumber (cm-1)

4

3800

3600

3400

3200

3000

3900

3600

Wavenumber (cm-1)

3300

3000

2700

2400

Wavenumber (cm-1)

Fig. 3. Polarized FTIR spectra of H-annealed (a) olivine, (b) clinopyroxene, and (c) plagioclase. The spectra were normalized to 1 cm thickness and vertically offset for clarity. X, Y and Z denote three mutually perpendicular directions (see text).

The annealed crystals showed no zonation in H-species. Nevertheless, some of the recovered samples after conductivity runs, e.g., some olivines in which H diffusivity is faster and the H-rich clinopyroxene /(100), demonstrated ~10–30% lower H-related IR absorbance in the outmost ~30–50 μm rims relative to the nearly constant absorbance in other regions (Fig. 4), indicating slight dehydration. However, most samples showed no change in water content during the runs, and even for the slightly dehydrated samples, the water contents measured out of the outmost rims were essentially the same as those initially annealed (Table 2). Representative complex spectra are shown in Fig. 5. Usually the spectra show an arc over the whole frequency range at low temperatures, and a high frequency arc plus a low frequency short tail at high temperatures. The arc in the high frequency range is due to sample interior conduction, and the tail is usually attributed to electrode effects (Huebner and Dillenburg, 1995). At low temperatures when the sample resistance is relatively large, the spectra show very much data scatter in the low frequency range (Fig. 5b; the scatter is also obvious in the so-called Bode-diagrams in terms of impedance vs. frequency and phase angle vs. frequency), probably due to increased impedances (Poe et al., 2010); while at high temperatures, the spectra in the low frequency range are usually related to electrode effects. Therefore, the measured impedance at a single low frequency,

e.g., 0.01 Hz in some reports (Yoshino et al., 2006), may be physically meaningless in reflecting the sample resistance. Electrical conductivity of silicate minerals as a function of temperature is usually expressed by the Arrhenius relation (given that oxygen fugacity and other factors remain constant):   ΔH σ ¼ A0 exp − RT

ð2Þ

where A0 and ΔH are the pre-exponential factor and activation enthalpy, respectively, R is the gas constant and T is temperature. Parameters from fitting Eq. (2) to the data are summarized in Table 2, and the measured data from different cycles and their fitting to Eq. (2) are illustrated in Figs. 6, 7 and 8 for olivine, clinopyroxene and plagioclase, respectively. In the first cooling cycle of the H-rich clinopyroxene (100), the conductivity is higher than that from the first heating cycle (Fig. 7a), and has a very small activation enthalpy of ~37 ± 2 kJ/mol similar to that of conduction by ionic water (Roberts, 2002), indicating partial dehydration at the highest temperature. This agrees with the FTIR profile analysis that the outmost rims dehydrated slightly (Fig. 4), and the related data were excluded in our analysis. The reproducible conductivity between different heating/cooling cycles (Figs. 6, 7 and 8) suggests the absence of system hysteresis, and implies that water content did not 1500

110

a

b 1400

Integral absorbance

Absorbance

100

90

80

1300

1200 70

60 4000

3800

3600

3400

3200

3000

1100

0

300

600

900

Wavenumber (cm-1) Fig. 4. (a) Unpolarized FTIR spectra and (b) their integral absorbance for a profile analysis on the recovered rim-dehydrated clinopyroxene /(100). The spectra were normalized to 1 cm thickness and vertically offset for clarity. Note that the dehydrated rims are so small (~30–50 μm) that polarized measurements on different directions are difficult, if not impossible.

X. Yang / Earth and Planetary Science Letters 317–318 (2012) 241–250

a

245

change a lot during the runs. The results show negligible anisotropy in electrical conductivity for hydrous olivine (Fig. 6d) and clinopyroxene (Fig. 7d) but relatively large anisotropy, by a factor of ~3–8, for hydrous plagioclase (Fig. 8d).

direction /(010) is much more conductive than along other directions. In their experiments, however, they increased temperature to ~700 °C first for ~1 h and then measured the complex spectra with decreasing temperature, in which the H-rich samples may have dehydrated at the highest temperatures. At their run conditions (8 GPa and b~700 °C for hydrous samples), the released water should be in the form of liquid or supercritical liquid according to the phase diagram of water (Dunaeva et al., 2010), and if such water did not escape immediately and completely, the subsequent conductivity measurements may have been greatly affected. In fact, the H-rich samples in Poe et al. (2010), with H2O contents >1000 ppm, generally had very small activation enthalpies, ~36–38 kJ/mol essentially the same as those of partially dehydrated samples during cooling cycles, e.g., clinopyroxene in this study and olivine in Yoshino et al. (2009). Therefore, the measurements of the H-rich olivines in Poe et al. (2010) may have been flawed by dehydration, and only their data from H-poor samples, with similar activation enthalpies as those by H conduction in olivines (see next section), are probably physically meaningful. The data from low water content olivines of Poe et al. (2010) are compiled in Fig. 6d. The samples with similar H2O contents, e.g., 585 ppm /(010) vs. 592 ppm /(100) and 363 ppm /(100) vs. 393 ppm /(001), show negligible anisotropy in conductivity, consistent with our data for the olivines with ~ 40 ppm H2O (Fig. 6d). The increased conductivity with increasing water content from these data agrees with early reports for hydrous olivine and other minerals (Dai and Karato, 2009a,b,c; Huang et al., 2005; Poe et al., 2010; Wang et al., 2006; Yang and McCammon, 2011; Yang et al., 2011a,b; Yoshino et al., 2009). Our results also show similar conductivity along different orientations for hydrous clinopyroxene (Fig. 7d). So far no published data are available for orientation-related conductivity of hydrous clinopyroxene, but Dai and Karato (2009a) reported negligible anisotropy for hydrous orthopyroxene. The small anisotropy of hydrous olivine and pyroxenes differs strongly from the large anisotropy of hydrous plagioclase (Fig. 8d). With similar composition in major elements between the starting materials (but very different Fe 3 +/Fetotal ratios), the conductivity of the single crystal clinopyroxenes with ~680 ppm H2O is lower than that of hydrous polycrystalline clinopyroxenes with less H2O reported by Yang and McCammon (2011) (Fig. 7d). This is very likely due to the different Fe 3 +/Fetotal ratios between the H-annealed samples, e.g., ~ 6 ± 4% in this study vs. ~ 34 ± 3% in Yang and McCammon (2011), as argued by Yang and McCammon (2011) that, given a water content (either dry or hydrous), a variation of ~10% Fe 3 +/Fetotal leads to a variation of conductivity by a factor of ~2–7 at 1300–400 °C.

4. Comparison with previous studies

5. Conduction mechanism

Yoshino et al. (2006) measured the electrical conductivity of hydrous olivines along different orientations, however, their results may have been biased by: (1) unpolarized FTIR method with the calibration of Paterson (1982), which may overestimate or underestimate water contents as argued by Bell et al. (2003), (2) single frequency conductivity analysis at 0.01 Hz, which may affect their data as noted above, (3) two of their three samples dehydrated largely in their runs, which makes it difficult to evaluate their data, and (4) the use of Ni foils as electrodes and oxygen buffer may introduce additional uncertainty (Appendix B). Hence, the arguments by Yoshino et al. (2006) that the conductivity of hydrous olivine is anisotropic, with direction /(100) the highest at low temperatures and direction /(001) the highest at high temperatures, must be treated with caution, and a direct comparison between their results and ours is not favored. Poe et al. (2010) determined the electrical conductivity of hydrous olivines along different orientations, with up to >2000 ppm H2O, and reported large anisotropy at elevated water contents that the

The electrical conductivity of a H-bearing silicate mineral can be further quantified as:

Z” (kOhm)

-15000

-10000

-5000

0 0

5000

10000

15000

20000

60000

80000

Z’ (kOhm)

b

Z” (kOhm)

-60000

-40000

-20000

0

0

20000

40000

Z’ (kOhm) Fig. 5. Representative complex spectra for (a) clinopyroxene /(100) and (b) olivine /(100). Z′ and Z″ are the real and imaginary parts of complex impedance, respectively. Frequency decreases from left (106 Hz) to right (20–0.1 Hz) along the Z′-axis. An equivalent circuit of a single R-CPE circuit element, a resistor R and a constant phase element (CPE) in parallel, was used to fit the impedance arc in the high frequency region; the short tail in the low frequency range and the offset from the origin were not included. Solid curves show the fit of the high frequency arcs. Inset shows the spectrum at (a) 600 °C and (b) 700 °C.

σ ¼ σD þ σH

ð3Þ

where σD and σH are the conductivities due to non-H and H-related defects, respectively. σD corresponds to the conductivity under dry conditions, and is usually attributed to small polarons in Fe-bearing minerals (e.g., electrons hopping between Fe 2 + and Fe 3 +) and to other charge carriers in Fe-free minerals (e.g., Na + in plagioclase). σD is usually smaller than σH, it is thus widely accepted that the charge transfer of hydrous minerals is dominated by H-related defects, including its direct effect due to the H movement and indirect effect due to the H-enhanced mobility of other species (Dai and Karato, 2009a,b,c; Huang et al., 2005; Poe et al., 2010; Wang et al., 2006; Yang and McCammon, 2011; Yang et al., 2011a,b; Yoshino et al., 2009). The activation enthalpies of the hydrous minerals (Table 2) differ from those reported for the dry polycrystalline/single

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X. Yang / Earth and Planetary Science Letters 317–318 (2012) 241–250

a

b

c

d

Fig. 6. Electrical conductivity of hydrous olivine (a) /(100), (b) /(010), (c) /(001), and (d) the linear fitting to all the data from (a), (b) and (c), respectively. Data of low water content San Carlos olivines from Poe et al. (2010), measured at 8 GPa, were shown for comparison. The solid gray line in (a) is the assembly background conductivity by replacing the sample with a piece of dense ceramic Al2O3 (Yang et al., 2011b), which increases quickly above ~ 950 °C. The high conductivity values measured at the highest temperatures, especially (b) and (c), were probably caused by a higher contribution from the assembly background conduction and/or sample dehydration at these temperatures. Shaded region illustrates the electrical conductivity of 0.01–0.1 S/m (temperature is assumed 1000–1300 °C for conductive zones in the uppermost mantle).

a

b

c

d

Fig. 7. Electrical conductivity of hydrous clinopyroxene (a) /(100), (b) /(010), (c) /(001), and (d) the linear fitting to all the data from (a), (b) and (c), respectively. Data measured on polycrystalline clinopyroxenes with different water contents from Yang and McCammon (2011), with similar composition in major elements between the starting materials (but with different Fe3 + contents: see text), were shown for comparison. Shaded region illustrates the electrical conductivity of 0.01–0.1 S/m (temperature is assumed 1000–1300 °C for conductive zones in the uppermost mantle).

X. Yang / Earth and Planetary Science Letters 317–318 (2012) 241–250

247

a

b

c

d

Fig. 8. Electrical conductivity of hydrous plagioclase (a) /(100)′, (b) /(010), (c) /(001)′, and (d) the linear fitting to all the data from (a), (b) and (c), respectively. Data measured on polycrystalline plagioclases with different water contents from Yang et al. (2011a), with similar composition in major elements between the starting materials, were shown for comparison. Shaded region illustrates the electrical conductivity of 0.0001–0.1 S/m (temperature is assumed 700–1000 °C for conductive zones in the lower continental crust).

crystal ones (even of similar composition in major elements for each), e.g., 84 ± 6 to 92 ± 7 kJ/mol vs. 110–180 kJ/mol for conduction by small polarons in dry olivine (Poe et al., 2010; Xu et al., 2000), 62 ± 1 to 65 ± 1 kJ/mol vs. 87–102 kJ/mol for conduction by small polarons in dry clinopyroxene (Yang and McCammon, 2011; Yang et al., 2011b), and 115 ± 5 to 130 ± 4 kJ/mol vs. 161 ± 6 kJ/ mol for conduction by Na + in dry plagioclase (Yang et al., 2011a). Accordingly, the conduction mechanism in our samples is controlled by H-species. Some recent well-conducted experiments have shown that the activation enthalpy of H-dominated conduction in hydrous minerals is independent of water content (Dai and Karato, 2009a,c; Romano et al., 2009; Yang and McCammon, 2011; Yang et al., 2011a,b; Yoshino et al., 2009). The activation enthalpies of olivines in this study are similar to those reported for hydrous polycrystalline (Yoshino et al., 2009) and single crystal olivines (Poe et al., 2010: with low H2O contents), e.g., 80–99 kJ/mol due to H conduction. The activation enthalpies of clinopyroxenes in this study resemble those reported for hydrous polycrystalline and single crystal clinopyroxenes (Yang and McCammon, 2011; Yang et al., 2011b), e.g., 67–74 kJ/mol due to H conduction. However, the activation enthalpies of plagioclases in this study differ from those reported for polycrystalline plagioclases (Yang et al., 2011a), e.g., 77 ± 2 kJ/mol due to H conduction. Water dissolved in the H-annealed olivines and clinopyroxenes in both this study and previous reports is all OH; by contrast, only OH is dissolved in the plagioclases of this study (no sharp bands above 3400 cm − 1 in the IR spectra which are typical of molecular H2O in plagioclase: Johnson and Rossman,

2004), but both OH and molecular H2O are present in the samples of Yang et al. (2011a), either separated from a xenolith granulite equilibrated at 10 kbar and ~ 900 °C or H-annealed with the same material at 10 kbar and 800 °C. We suggest that the different activation enthalpies of the hydrous plagioclases between this study and Yang et al. (2011a) are caused by the different H-species. This may also explain the different conductivity between these samples given similar H2O contents (~ 105 vs. 135 ppm: Fig. 8d). The activation enthalpies of H conduction in olivine, clinopyroxene and plagioclase are generally smaller than those measured from H diffusion, e.g., ~ 110 to 200 kJ/mol for olivine, ~ 110 to 180 kJ/mol for clinopyroxene and ~170 to 220 kJ/mol for alkali feldspar and plagioclase (Ingrin and Blanchard, 2006); also, diffusion experiments have shown that H diffusivity is anisotropic in olivine and pyroxenes but probably isotropic in plagioclase (Ingrin and Blanchard, 2006), different from H conductivity in these minerals (Figs. 6d, 7d and 8d). Given H conduction and diffusion due to intrinsic properties of silicates, these differences may be related to: (1) a concentration gradient, and thus a higher chemical potential gradient, is involved in the measurement of activation enthalpy in diffusion but not in conduction experiments; (2) because of the concentration gradient, the dissolved H-species are not in equilibrium in diffusion but in equilibrium in conduction experiments; (3) different charge carriers operate mutually in hydrous minerals as indicated by Eq. (3); and (4) based on (1) and (2), H mobility may be different between conduction and diffusion experiments, and the presence of different H defects, e.g., OH vs. molecular H2O, may strongly affect conduction as mentioned above for plagioclase (Fig. 8d).

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The above discussions demonstrate that the mobility of H is not necessarily the same between conduction and diffusion, and that the activation enthalpy is usually smaller in conduction than in diffusion experiments. Recently, Yang et al. (2011a) also found that, for dry plagioclase, the activation enthalpy of Na transfer seems smaller in conduction than in diffusion. Therefore, it is not appropriate to estimate electrical conductivity by theoretical modeling using the Nernst–Einstein relation with diffusivity of species from diffusion experiments via: 2

σ¼

f Dcq RT

ð4Þ

where f is a constant, D, c and q are the self-diffusion coefficient, concentration and electrical charge of the charged species, respectively. 6. Geophysical implications Available experiments have shown that, given water content not changed a lot, H conduction continues up to ~ 1100 °C for pyroxenes, garnet and wadsleyite (Dai and Karato, 2009a,b,c; Romano et al., 2009; Yang and McCammon, 2011; Yang et al., 2011a,b). We thus assume that our data by H conduction can be extrapolated to higher temperatures as generally assumed in previous reports (Dai and Karato, 2009a,b,c; Huang et al., 2005; Poe et al., 2010; Romano et al., 2009; Wang et al., 2006; Yang and McCammon, 2011; Yang et al., 2011a,b; Yoshino et al., 2009), although our measurements were only made up to 700–800 °C. In combination with approaches from petrology, geochemistry, geophysics and laboratory studies, our data provide important constraints on the electrical anisotropy and structure of the lower crust and uppermost mantle. 6.1. Electrical anisotropy in the lower continental crust The main constituents of the lower crust are mafic granulites dominated by plagioclase, clinopyroxene and orthopyroxene, of which the chemical and modal compositions are highly variable from one place to another and even from one sample to another, and the prevailing temperature from xenolith data, seismic imaging and heat flow is mostly ~ 700–1000 °C (Rudnick and Fountain, 1995; Rudnick and Gao, 2003). Because temperatures above the solidus are only occasionally reached in the crust (and thus melt, if present, should be very limited), high electrical conductivity in the lower crust (mostly 0.0001–0.1 S/m), probably on a global scale with some local exceptions, was usually attributed to inter-connected grain boundary aqueous fluids or graphite films (Jones, 1992) (and thus, seismic reflections in the lower crust were interpreted in some reports by the presence of fluids/solidified intrusions (Mooney and Meissner, 1992)). However, each of these explanations may be only locally important, and is not expected as a general mechanism which has been discussed in detail elsewhere (Hyndman et al., 1993; Jones, 1992; Yardley and Valley, 1997). In brief, the presence of up to several percent of aqueous fluids, as required by the high conductivity, is difficult to reconcile with the physics of a fluid–solid mixture (in which the effective compaction should drive the fluids out of the system if they are different in density from the solid), the high-grade metamorphic lithology of the lower crust should exhaust all the hydrous fluids via retrograde reactions, and the highly heterogeneous composition of some fluid-sensitive trace elements and Li–C–O isotopes in the lower crust on both large and small scales are incompatible with a widespread hydrous fluids; also, no known mechanism can lead to a general connection of graphite films in the lower crust and their disconnection at, for example, ~ 20/40 km depth in average so that the conductivity above/below the lower crust is markedly decreased. Moreover, conduction by these highly conductive phases has very small temperature dependence and is difficult to explain the more conductive lower crust

in Phanerozoic than in Precambrian areas, e.g., 0.03–0.1 S/m in the former vs. 0.003–0.01 S/m in the latter for an assumed layer of 10 km thickness (Hyndman and Shearer, 1989), which is probably related to a temperature contrast between these regions, by ~ 100–300 °C, due to thicker lithospheric roots and thus poor thermal conduction and lower temperature in the latter (Artemieva, 2009). Pyroxenes in the lower crust are rich in Fe (and Fe 3 +), e.g., by a factor of ~2 to >7, than those in mantle peridotites, and plagioclase in the lower crust generally contains higher amounts of Na2O, e.g., ~3–7%, which species may act as important charger carriers. Recently, Yang et al. (2011a,b) measured the electrical conductivity of plagioclase, clinopyroxene and orthopyroxene prepared from mafic xenolith granulites under well-controlled lower crustal conditions, as well as the effect of H on conduction, and suggested that the high conductivity in the lower crust and the lateral heterogeneity can be mostly accounted for by the main constitutive minerals themselves due to chemical composition (Fe 3 +, Na + and H2O) and temperature, without significant contributions from other secondary phases. Accepting the results of Yang et al. (2011a,b), the resolved electrical anisotropy may originate from the fabrics of the main constituents: (1) LPO of hydrous plagioclase. Due to the negligible conductivity anisotropy of hydrous pyroxenes and large conductivity anisotropy of hydrous plagioclase (Dai and Karato, 2009a; and this study), hydrous plagioclase may lead to electrical anisotropy where it is abundant and has a strong LPO, e.g., due to ductile deformation in the lower crust; and (2) microstructure of granulites. Yang et al. (2011a,b) found that, given equilibrium partitioning of elements (under either dry or hydrous conditions), the electrical conductivity of lower crustal clinopyroxene and orthopyroxene is similar but higher by a factor of ~ 4 to 20 than that of plagioclase, depending on composition and temperature. This, along with the lithology that mafic xenolith granulites usually show alternating pyroxene- and plagioclase-rich bands/layers on mm to cm scales (Yang et al., 2011b), will produce substantial electrical anisotropy. The banded fabrics of lower crustal minerals and their potential LPO may also lead to seismic reflectivity (e.g., Mooney and Meissner, 1992), because of the different seismic properties between pyroxenes and plagioclase. In this case, electrical anomalies and seismic reflectivity in the lower crust can be accounted for by the main constituents. This model can also explain why conductive zones and seismic reflectivity are limited within the lower crust but absent above and below in many areas (Jones, 1992; Mooney and Meissner, 1992), because the constitutive minerals, compositions, fabrics and temperatures are very different between the lower crust and the overlying shallow crust/underlying upper mantle. 6.2. Electrical anisotropy in the uppermost mantle 6.2.1. Grain boundary materials and hydrous olivine In contrast to the lower continental crust where conductive zones are common, only some regions of the uppermost mantle are of electrical anomalies. For example, low conductivity down to ~ 0.0003– 0.001 S/m has been reported for some areas in the upper mantle, e.g., at depths of ~40–180 km beneath Western Canada (Boerner et al., 1999). Because the conductivity of dry olivine is b0.001– 0.01 S/m at mantle conditions (Constable, 2006), electrical anomalies in the uppermost mantle, ~0.01–0.1 S/m, have been often attributed to inter-connected grain boundary partial melt (Gaillard et al., 2008; Shankland et al., 1981; Yoshino et al., 2010), graphite (Duba and Shankland, 1982) or sulfides (Watson et al., 2010). Melt at depths of ~40–200 km, if present, is mainly silicic in composition, and the fractions are usually b0.1–0.3% (Hirschmann, 2010); by contrast, up to several percent of such melt is required for the mantle conductive zones (Shankland et al., 1981; Yoshino et al., 2010). Also, petrological and geochemical studies of natural samples have shown that modal/ cryptic melt metasomatism is quite common in the uppermost mantle (Menzies et al., 1987; Roden and Murthy, 1985), indicating

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substantial permeability. Accordingly, maintaining a continued presence of a large amount of melt over a broad region in the uppermost mantle is physically difficult. On the other hand, it is not clear if and how inter-connected graphite/sulfides are present on appreciated scales in the relatively oxidizing uppermost mantle, where the oxygen fugacity is usually ±2 log units relative to the fayalite– magnetite–quartz (FMQ) buffer and is beyond the stability field of graphite and most sulfides (Frost and McCammon, 2008). Moreover, recent studies on partial molten systems demonstrate that a network of grain boundary melt, either in pervasive forms under static conditions (Zhu et al., 2011) or in planar melt-rich tube forms under stressrelated active conditions (Kohlstedt and Holtzman, 2009), cannot lead to the observed electrical anisotropy in some regions of the upper mantle, e.g., the southern East Pacific Rise (Evans et al., 2005); also, aligned distribution of inter-connected grain boundary graphite or sulfides only along certain direction(s) has so far not been reported for mantle samples (e.g., xenolith peridotites). Therefore, conduction by any of these materials may play roles only in regional zones, but is not expected as a general mechanism for all the resolved electrical anomalies. Some scholars argued that the conductive zones may be due to H in hydrous olivine (Karato, 1990; Wang et al., 2006). However, olivines in the uppermost mantle usually contain b100 ppm H2O (Appendix C), and the electrical conductivity is only ~ 10 − 3 S/m at 1200 °C for olivine with ~ 40 ppm H2O by extrapolating our data (Fig. 6d; note that, based on the combined data by Yang et al. (2011a), the conr ductivity of hydrous minerals is proportional to water content, Cw , with the exponent r ≈ 1, and a change of ~ 40 to 100 ppm H2O in olivine may only cause a change of conductivity by a factor of ~2). Similar low conductivity of H-poor olivines has also been reported by Yoshino et al. (2009) and Poe et al. (2010). Moreover, the nearly isotropic conductivity of hydrous olivine and pyroxenes (Dai and Karato, 2009a; and this study) indicates that they cannot lead to electrical anisotropy. In addition, xenolith peridotites usually show no banded mineral fabrics (Mercier and Nicolas, 1975), excluding the possibility that electrical anisotropy is related to the fabrics of minerals due to different conductivities (as argued above for the lower crust). Therefore, hydrous olivine and pyroxenes in peridotites cannot produce the observed electrical anomalies. 6.2.2. Macro-heterogeneity in the uppermost mantle The uppermost mantle is dominated by peridotites consisting of olivine, orthopyroxene and clinopyroxene. Due to profound diversity in petrology and geochemistry, however, the upper mantle is macroscopically heterogeneous, and at least two other materials may be regionally important. First, augites and other pyroxenites in the form of megacrysts, xenocrysts and xenoliths are by far the most abundant deep-seated samples captured by volcanic-eruptions only less than coexisting xenolith peridotites in a wide variety of tectonic settings, and may have crucial influence on basalt genesis and the structure and refertilization of the upper mantle (Allègre and Turcotte, 1986; Hirschmann et al., 2003). These materials can be present as dykes/ veins at depths of ~40–200 km (Davies et al., 2001; Lehtonen et al., 2009; Shaw and Eyzaguirre, 2000). Second, the enriched composition in incompatible elements of many areas in the upper mantle, as observed for mid-ocean ridge basalts (MORB), oceanic-island basalts (OIB) and mantle xenoliths, indicates that subduction-related former oceanic crust and possibly delaminated lower continental crust were involved in the evolution of the uppermost mantle (Allègre and Turcotte, 1986; Hofmann, 1997). Remnants of the recycled materials may have a survival time of ~20 to >150 Ma in the upper mantle, without being melted or homogenized, and become deformed, folded, thinned and stretched into high aspect ratios lenses by mantle flow (Allègre and Turcotte, 1986; Anderson, 2006). Usually, augites and other pyroxenites contain more Fe (and Fe 3 +) than pyroxenes in mantle peridotites, e.g., ~ 5–12 wt.%

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FeOtotal and ~25–45% Fe 3 +/Fetotal in augites (Bell et al., 2004; Dyar et al., 1996; Shaw and Eyzaguirre, 2000) vs. ~2–3 wt.% FeOtotal and ~10–30% Fe3 +/Fetotal in peridotite diopsides (Bell et al., 2004; Canil et al., 1994; Dyar et al., 1996; Shaw and Eyzaguirre, 2000), and pyroxenes contain more water than olivine, e.g., up to >1000 ppm H2O in the former (Bell and Rossman, 1992; Ingrin and Skogby, 2000). H-rich augite has high conductivity of 0.01–0.1 S/m at mantle temperatures, and the conductivity can be further enhanced by increasing Fe3 + content (Fig. 7d: Yang and McCammon, 2011; and this study). Accordingly, a local presence of moderate Fe3 +- and H2O-rich augites and other pyroxenite analogs, e.g., b10% depending on composition, temperature and modeling, can result in a bulk conductivity of ~0.01– 0.1 S/m (Yang and McCammon, 2011). On the other hand, the remnants of recycled crustal materials may be conductive, due to enrichedvolatiles and higher concentrations of Fe3 + and other charged species (e.g., gabbros in subducted oceanic crust and granulites in delaminted continental crust: Yang et al., 2011a,b), and they can also lead to a bulk conductivity of ~0.01–0.1 S/m due to a local enrichment, e.g., b20% depending on composition, temperature and modeling given a similar conductivity as the granulite minerals in Yang et al., 2011a,b. Therefore, both these components have the potential yielding high conductivity in the uppermost mantle. The spatial relation of these materials relative to wall peridotites, e.g., dykes/veins of augites and other pyroxenites and/or lenses of remnants from recycled crustal materials controlled by mantle flow, can lead to substantial electrical anisotropy. For example, the conductivity is higher in the direction of the augiteand pyroxenite-dykes/veins or remnant-lenses than along other directions, because the wall peridotites, dominated by olivines, are less conductive. In summary, electrical anomalies in the uppermost mantle are probably not caused by hydrous olivine (and pyroxenes in mantle peridotites). In addition to the potential roles played by inter-connected grain boundary partial melts and/or other conductive materials, a local presence of minor amounts of Fe 3 +- and H2O-rich augites and other pyroxenites and/or remnants of recycled crustal materials, along with their spatially heterogeneous distribution, may be responsible for the electrical anisotropy in regional uppermost mantle. The local enrichment of such materials would reduce seismic velocity, leading to low-velocity zones (LVZ) (Fischer et al., 2010); moreover, some other chemical-physical properties of the uppermost mantle may also be affected, such as melting, heterogeneity, rheology and anelasticity. The suggested model links mantle petrology, geochemistry and geophysics, and can reconcile some conflicting arguments on some properties of the upper mantle, e.g., arising from approaches of different Earth Science disciplines. Future work on the conductivity of other pyroxenites and possible recycled materials and on their spatial distribution by laboratory experiments and numerical modeling may provide new constraints on the electrical properties and structure of the uppermost mantle.

Acknowledgments XY is thankful to Andreas Audétat for supplying the starting olivines, and to Li Zhang, Hubert Schulze, Julia Huber, Catherine McCammon, Heinz Fischer, Stefan Übelhack, Hongzhan Fei and Yuan Li for technical assistance. Constructive comments from two anonymous reviewers greatly improved the manuscript. This work was partly supported by the Natural Science Foundation of China (40903016) and the Alexander von Humboldt Foundation.

Appendix A. Supplementary data Supplementary data to this article can be found online at doi:10. 1016/j.epsl.2011.11.011.

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