Effect of fluid H₂O on compressional wave velocities in quartz aggregate up to 500°C at 0.5 GPa

The high-pressure apparatus and layout of the experimental system for conducting the ultrasonic experiments in this study are shown in Figure 1. The elements of this experimental setup included a piston cylinder press and an ultrasonic system. We used a solid-medium piston cylinder apparatus with an internal diameter of 34 mm and a length of 80 mm (Toshiba Tungaloy Company, Kawasaki City, Japan). Talc, pyrophyllite, and boron nitride (BN) were used as the solid pressure transmitting medium that also contained a heating element and sample. The calibration of pressure has been described elsewhere (Kono et al. 2004,2008; Kozai and Arima 2005; Nishimoto et al. 2005).

Both top and bottom faces of the core sample (approximately 6.0 mm diameter and 6.0 mm length) were diamond-polished. Each core sample was vacuum-dried at 180°C for 24 h. The vacuum-dried samples were then encapsulated in platinum with or without distilled water. A 20-mm long platinum buffer rod was placed between the platinum-encapsulated sample and a LiNbO₃ piezoelectric transducer. The piezoelectric transducer was mounted on the bottom of the platinum buffer rod. No glue was used at any interface. The encapsulated sample, buffer rod, and piezoelectric transducer were then placed in the talc pyrophyllite high-pressure cell. High temperatures were achieved using a graphite heater and were monitored by a Pt-PtRh₁₃ thermocouple. A thermocouple junction was placed on the top of the platinum-encapsulated sample. The temperature differences between the thermocouple junction and several parts of the cylindrical rock sample were measured to evaluate the thermal distribution of the cell assembly (Kono et al. 2004,2008). Figure 1 also shows details of the high-pressure, high-temperature cell assembly inside the cylinder.

Ultrasonic velocity measurements were performed using the pulse reflection technique with buffer rod. To generate elastic waves, sinusoidal electric burst signals were produced by an arbitrary waveform generator (Tektronix AWG2021, Beaverton, OR, USA) and applied to the LiNbO₃ piezoelectric transducer (10° Y-cut). The elastic waves of 8 MHz were transmitted through the buffer rod and the sample. The sample length was more than ten times longer than the wavelength of the elastic waves. Reflected elastic waves from the top and bottom surfaces of the sample were received by the transducer and then converted into an electric signal. The waveforms were recorded using a digital oscilloscope (Hewlett-Packard 54825A, Palo Alto, CA, USA). These signals were sampled 4,096 times for each run at a sample rate of 1.0 × 10⁹ samples/s for a given P-T condition. We used the mean value of the recorded travel time for a given P-T condition and corresponding ray path to compute the average velocity. The maximum uncertainty in the measured V p values was ±0.4%. This procedure is described in more detail elsewhere (Kono et al. 2004).

The rock sample used in this study was a natural quartz aggregate (fine-grained silicified rock with a density of 2.58 g/cm³ and porosity of 3.2 vol.%) that was collected from the Iwakuni-Yanai area in the Ryoke belt, SW Japan (Yamamoto et al. 2004; Matsumoto et al. 2010). The Ryoke belt in this area is mainly composed of a metamorphosed Jurassic accretionary complex and Cretaceous granites (Ikeda 1998). Fine-grained silicified rocks occur in pelitic schist of the biotite zone and in the northernmost part of the cordierite zone (Yamamoto et al. 2004). The sample used in this study was composed mainly of quartz (about 90 vol.%) with minor muscovite, biotite, feldspar, and opaque minerals (Figure 2), which were <50 μm in grain size. The rock shows a granoblastic and nearly isotropic texture.

The measurements were carried out up to 0.5 GPa (wet samples) or 0.6 GPa (dry samples) at temperatures ranging up to 500°C using a piston cylinder apparatus and the techniques described above. In this experiment, the pressure was first increased to 0.5 GPa (wet samples) or 0.6 GPa (dry samples) at room temperature and then the temperature was increased to 500°C. To ensure a steady state travel time, we kept the P-T conditions of the experiment constant for 30 to 180 min before performing the velocity measurements. The sample recovered elastically and retained its original length, which suggests that it had not been subjected to plastic deformation during the entire experiment. The compressional wave velocities are plotted against pressure in Figure 3 and against temperature in Figure 4.

As shown in Figure 3, the compressional velocities in the dry sample increased rapidly until a pressure of approximately 0.3 GPa was reached. The increase in V p was probably attributable to the closure of microcracks (e.g., Birch 1960; Kern 1982). The compressional wave velocities also showed a linear increase at higher pressure between 0.3 and 0.6 GPa, which suggests that the effect of porosity on the measured velocities was not significant within this pressure range. By fitting velocity data using linear functions of pressure, we obtained the pressure derivative ∂V p/∂P = 0.29 km s^-1 GPa^-1. The pressure derivative (in the range P = 0.3 to 0.6 GPa) was then used to linearly extrapolate the velocity data, which yielded a V p value of 6.04 km s^-1 under ambient conditions. This is the so-called intrinsic velocity of Kern (2011). As in previous studies (Shingai et al. 2001; Kitamura et al. 2003; Kono et al. 2004; Nishimoto et al. 2005; Ishikawa et al. 2008), linear trends of velocity as a function of pressure were observed above a threshold pressure of between 0.3 and 0.6 GPa. At lower pressures, the velocities of the samples are compromised by unclosed pore spaces filled with air.

The dry sample was heated from room temperature at 0.5 GPa to 500°C. Unfortunately, the lead wire was disconnected from the piezoelectric transducer within the high-pressure cell at 500°C and therefore V p could not be measured at 500°C. As the temperature increased to 400°C, the compressional velocities decreased from 6.19 to 5.99 km s^-1 (Figure 4). By fitting the velocity data using linear functions of temperature, we obtained the temperature derivative ∂V p/∂T = -0.52 × 10^-3 km s^-1°C^-1. Because the compressibility and thermal expansion of low quartz is less than 0.5% between ambient conditions and 0.5 GPa and 400°C (e.g., Raz et al. 2002), the changes of sample length can be considered to be insignificant for the determination of V p.

To understand the effect of water on V p, we measured the compressional wave velocities in wet samples (Figures 3 and 4). In each experiment involving a wet sample, the pressure was first increased to 0.5 GPa at room temperature (Figure 3). The compressional velocities increased rapidly up to a pressure of 0.3 GPa. This is probably attributable to the closure of air-filled pore spaces. The compressional wave velocities also increased linearly at higher pressures between 0.3 and 0.5 GPa, and these values showed a similar pressure dependence for different H₂O contents, which suggests that the effect of pore air on the measured velocities is not significant within this pressure range. By fitting velocity data using linear functions of pressure, we obtained the pressure derivative ∂V p/∂P = 0.45 km s^-1 GPa^-1 for 0.37 wt.% H₂O and 0.54 km s^-1 GPa^-1 for 0.98 wt.% H₂O. As before, we used the pressure derivative to linearly extrapolate the velocity data (in the range P = 0.3 to 0.5 GPa) to ambient conditions, which resulted in an intrinsic V p of 5.79 km s^-1 for 0.37 wt.% H₂O and 5.65 km s^-1 for 0.98 wt.% H₂O. After pressurization to 0.5 GPa, the sample was heated to 500°C (Figure 4), which caused the compressional velocities of the quartz aggregate with 0.37 wt.% H₂O to decrease from 6.02 to 5.79 km s^-1. By fitting the velocity data using linear functions of temperature, we obtained the temperature derivative ∂V p/∂T = -0.49 × 10^-3 km s^-1°C^-1 for quartz aggregate with 0.37 wt.% H₂O. After repeating the experiment using quartz aggregate with 0.98 wt.% H₂O, the compressional velocity decreased from 5.92 to 5.78 km/s as the temperature increased to 500°C. By fitting the velocity data using linear functions of temperature, we obtained ∂V p/∂T = -0.28 × 10^-3 km s^-1°C^-1. Comparison of the temperature derivatives of wet samples with that of the dry sample showed that the temperature derivatives were insensitive to the addition of fluid H₂O within the 1 wt.% range investigated.

The velocity drop is plotted against temperature in Figure 5. We used the following definition of the velocity drop:

$\begin{array}{c}\mathrm{Velocity}\mathrm{drop}=100\times \left[\mathit{V}{\mathrm{p}}_{\left(\mathrm{dry},\mathit{P}=0.5\mathrm{GPa},\mathit{T}\right)}-\mathit{V}{\mathrm{p}}_{\left(\mathrm{wet},\mathit{P}=0.5\mathrm{GPa},\mathit{T}\right)}\right]/\hfill \\ \mathit{V}{\mathrm{p}}_{\left(\mathrm{dry},\mathit{P}=0.5\mathrm{GPa},\mathrm{T}\right)}\hfill \end{array}$

The velocity drop was approximately 3% in the experiment using quartz aggregate with 0.37 wt.% H₂O. In the experiment with higher H₂O content (0.98 wt.% H₂O), the velocity drop remained constant (4%) up to 200°C and then decreased slightly to 3% with increasing temperature.

Zhao et al. (1996) reported a low seismic velocity (-5%) and high Poisson's ratio (+6%) anomaly extending over approximately 300 km² at the hypocentral zone of the magnitude 7.2 Kobe earthquake that occurred on January 17, 1995. They speculated that this anomaly was due to an over-pressurized, fluid-filled, fractured rock matrix that contributed to the initiation of the Kobe earthquake. The low velocity anomaly (-5%) is almost equivalent to the velocity drop (-4%) shown in our experimental results with 1 wt.% H₂O. Our results also show that the addition of 0.4 to 1.0 wt.% H₂O into the quartz aggregate causes a velocity drop of 3% to 4% at the pressures and temperatures of the middle crust (0.5 GPa, 25°C to 500°C), which suggests that one possible reason for low seismic velocities (-5%) is 1 wt.% H₂O fluid-filled rock in the hypocentral zone. At this stage, we see our experimental data as a useful contribution towards improved quantitative understanding of the effect of H₂O on the elastic wave velocities of rocks. We also hope our results will be useful for determining the geofluid distribution within the crust, thereby enhancing future interpretations of seismic tomographic data.

A compressional wave velocity of V p = 6.04 km s^-1 for dry quartz aggregate under ambient conditions was determined by linear backward extrapolation of the velocity data recorded at high pressures (0.4 to 0.6 GPa). The pressure derivatives obtained by fitting the high pressure data of 0.4 to 0.6 GPa yielded ∂V p/∂P = 0.29 km s^-1 GPa^-1. The temperature derivatives obtained at a pressure of 0.5 GPa and temperatures from 25°C to 400°C yielded ∂V p/∂T = -0.52 × 10^-3 km s^-1 °C^-1.

The compressional wave velocities at ambient conditions were determined to be V p = 5.79 km s^-1 for wet quartz aggregate of 0.37 wt.% H₂O and V p = 5.65 km s^-1 for wet quartz aggregate of 0.98 wt.% H₂O. These values were determined using backward extrapolation from the velocity data at higher pressure (0.4 to 0.6 GPa). The pressure derivatives obtained by fitting the high pressure data of 0.4 to 0.6 GPa yielded ∂V p/∂P = 0.45 km s^-1 GPa^-1 for the 0.37 wt.% H₂O sample and 0.54 km s^-1 GPa^-1 for the 0.98 wt.% H₂O sample. The temperature derivatives obtained at a pressure of 0.5 GPa and temperatures from 25°C to 500°C yielded ∂V p/∂T = -0.49 × 10^-3 km s^-1°C^-1 for quartz aggregate with 0.37 wt.% H₂O and ∂V p/∂T = -0.28 × 10^-3 km s^-1°C^-1 for quartz aggregate with 0.98 wt.% H₂O.

Our results demonstrate that the magnitude of the velocity drop is dependent on the degree of saturation of the pore spaces with fluid H₂O. The velocity drop was approximately 3% with 0.37 wt.% H₂O and 3% to 4% with 0.98 wt.% H₂O. Hence, a small amount of pore fluid (0.4 to 1.0 wt.%), in this case, water, can significantly reduce the compressional wave velocity under the P–T conditions of the middle crust.