Permeability and elastic properties of rocks from the northern Hikurangi margin: Implications for slow-slip events
Nicola Tisato1, Carolyn D. Bland1*, Harm Van Avendonk2, Nathan Bangs2, Hector Garza1, Omar Alamoudi1, Kelly Olsen2†, Andrew Gase2‡
1 Department of Earth and Planetary Sciences, Jackson School of Geosciences, The University of Texas at Austin, Austin, TX, U.S.
2 Institute for Geophysics, Jackson School of Geosciences, The University of Texas at Austin, Austin, TX, U.S.
* Now at: Pariveda Solutions, Dallas, TX, U.S.
† Now at: Descartes Labs, 1607 Paseo de Peralta, Suite B, Sante Fe, NM, U.S.
‡ Now at: Geology Department, Western Washington University, Bellingham, WA, U.S.
Corresponding author: Nicola Tisato (nicola.tisato@jsg.utexas.edu)
Key Points:
· Elastic properties, plastic deformation, and permeability of northern Hikurangi margin rocks
· Permeability-porosity relationship in accretionary prisms
· Clay swelling and plastic deformation controls permeability healing, providing a mechanism justifying slow-slip event cyclicity
Abstract
Fluid flow and pore-pressure cycling are believed to control slow slip events (SSEs), such as those that frequently occur at the northern Hikurangi margin (HM) of New Zealand. To better understand fluid flow in the forearc system, we examined the relationship between elastic properties, compaction, porosity, and permeability of Cretaceous-to-Pliocene sedimentary rocks from the Raukumara peninsula. We found that the permeability of the deep wedge is too low to drain fluids, but fracturing increases permeability by orders of magnitude, making fracturing key for fluid flow. In weeks to months, plastic deformation and clay swelling heal the fractures, restoring the initial permeability. We conclude that overpressures at the northern HM might partly dissipate during SSEs due to enhanced permeability near faults. However, in the weeks to months following an SSE, healing in the prism will lower permeability, forcing pore pressure to rise and a new SSE to occur.
Plain Language Summary
Earth's crust is composed of many tectonic plates fitting together like jigsaw puzzle pieces. Tectonic plates subduct in the mantle along active converging margins, where the forces driving such a convergence can trigger large earthquakes. However, these subduction zones often deform without producing earthquakes, but through slow-slip. The Hikurangi Margin (HM) of New Zealand is a well-studied subduction zone, producing both earthquakes and slow-slip events. The northern HM exhibits more frequent and shallower slow-slip events than the southern margin. Understanding what controls such differences can help improve the general understanding of subduction zone fault mechanics and earthquakes. One of the hypotheses is that the differences between the deformation of the northern and southern HM are controlled by the pore pressure at depth. We tested the elastic and fluid-transport properties of four samples from the northern HM and found that the overriding plate, if not fractured, would be impermeable to fluids. We also tested a fractured sample and observed efficient healing that resets the initial permeability. We conclude that fracturing the overriding plate is fundamental to draining the fluids carried at depth by the subducting plate, and slow-slip events may create new pathways for fluids to escape to the seafloor.
1 Introduction
At the shallow (<15 km depth) portion of the plate interface of subduction zones, scientists have found that convergence between the tectonic plates is often accommodated by modes of slip in between fast earthquakes and aseismic creep (Saffer & Wallace, 2015). Slow-slip events (SSEs) represent one class of such transient phenomena, which can lead to several centimeters of slip over several days to months (Schwartz & Rokosky, 2007). The relatively large seismic moment released by shallow SSEs, comparable to that of earthquakes (Passarelli et al., 2021), proves the importance to understand SSEs and how they influence the seismogenic character of a convergent margin. Frictional properties and stress heterogeneities along the plate interface might favor SSEs (Barnes et al., 2020; Bell et al., 2010; Im et al., 2020; Rabinowitz et al., 2018). Subducting oceanic crust and sediments release large volumes of fluids (i.e., seawater and CO2) whose pressure can exceed hydrostatic conditions when confined within low permeability rocks, lowering the effective stress on the shallow megathrust or splay faults and creating conditions conducive to SSEs (Kitajima & Saffer, 2012; Tsuji et al., 2008; Warren-Smith et al., 2019).
The northern Hikurangi margin (HM) of North Island, New Zealand, is a subduction zone with a shallow forearc and plate interface, where sediment accretion, compaction, and deformation have been modulated for millions of years by underthrusting seamounts (Gase et al., 2021; Sun et al., 2020). Subducting topography (e.g., seamounts) may cause stress heterogeneities (Bangs et al., 2023; Leah et al., 2022; Sun et al., 2020) and fluid pressure transients (Shaddox & Schwartz, 2019) that can lead to SSEs, several of which have been characterized in great detail by onshore geodetic and offshore absolute pressure gauge (APG) data (Yohler et al., 2019). Offshore Gisborne SSEs occur every 1-2 years and can last several weeks, during which 5 to 30 cm of slip may be accommodated (Wallace, 2020). Temporal variations in the character of earthquake focal mechanisms within the subducting oceanic crust provide compelling evidence for low effective stress before an SSE (Warren-Smith et al., 2019). This observation suggests that increases in fluid pressure enable SSEs and that the slip itself is accompanied by fluid release. Nevertheless, fluid transport through the accretionary wedge in this deformation cycle is not yet well understood (Antriasian et al., 2018).
The physical properties of accreted sediments of the northern HM and their relationship to slip phenomena have been studied recently with the use of cores and data from IODP expeditions (e.g., Wallace et al., 2019). The resulting studies have shed new light on the frictional properties, shallow dewatering, and faulting near the seafloor (Aretusini et al., 2021; Boulton et al., 2019, 2022; Dutilleul et al., 2021; Fagereng et al., 2019; French & Morgan, 2020; Shreedharan et al., 2022). However, to understand how fluid flow and deformation interplay in the deeper prism, we also must consider the physical properties of older, compacted, and diagenetically mature strata (Bland et al., 2015, Bassett et al., 2022). Here we present and discuss laboratory testing performed on rock samples from the subaerial northern HM as proxies of deep rocks in the prism to better understand fluid transport within the subduction zone.
2 Materials and Method
To test the compaction, elastic, and transport properties of rocks from the northern HM, we collected and performed experiments on outcrop samples from the Raukumara peninsula (Figs 1, S1) presenting different ages and degrees of diagenesis. In the central part of the peninsula, we collected a fine-grained sandstone from the Jurassic-to-Early Cretaceous Torlesse Supergroup forming the backstop for the accretionary wedge (sample MO02) (Adams & Graham, 1996; Mortimer et al., 2014). Just east of sample MO02 location, we sampled a calcareous fine-grained sandstone with a silty matrix from the Late Cretaceous-to-Paleocene Tinui Group (sample MT07) that likely represents an early passive margin deposit, now deeply buried in the accretionary wedge (Mortimer et al., 2014). Closer to the East coast, we collected a siltstone (sample GB13) from the middle Miocene Tolaga Group, which was deposited in slope basins after subduction initiated along the HM (van de Lagemaat et al., 2022), and a glauconitic fine-grained sandstone (sample FB12) from the Pliocene Mangaheia Group.
We determined mineral abundances and assemblages of each sample through X-ray diffraction (XRD) analyses and transmitted light microscopy by preparing 30 mm in thickness thin-sections. To estimate density, porosity, compressional and shear ultrasonic wave velocities (i.e., Vp and Vs), and helium gas permeability, we prepared cylindrical core plugs with parallel end faces for each sample. Samples were tested at the UT Austin Rock-Deformation-Laboratory for confining pressures (Pc) up to 200 MPa (~12.5 km depth for hydrostatic pore pressure and overburden density of 2.6 g/cm3) and deviatoric vertical force (Fv) ~2.6 kN. Each core plug was mounted inside a PVC jacket and between two core holders equipped with ultrasonic transducers and fluid ports, which are used to saturate and measure the core permeability. This sample assembly is mounted inside the triaxial cell (NER Autolab 1500) between the load cell and the vertical force piston. We define the mean stress as , where s2 = s3 = Pc and s1 is the maximum vertical stress: , where is the deviatoric stress, and A is the sectional area of the core plug. We also define effective stress ( ) as the difference between the mean stress and the pore pressure: .
We measured ultrasonic velocities using the transmission method at room temperature and a frequency of ~800 kHz (Birch, 1960). To understand the effect of saturation on Vp and Vs, we measured the ultrasonic velocities of sample GB13 saturated with water previously chemically equilibrated with the sample. During 30 hours, we recorded the injection of 4.7 ml of this fluid, equivalent to 136% of GB13 pore-space volume.
Sample permeabilities were calculated through the transient method measuring the pressure equilibration of the helium gas contained in two volumes connected to the sample end-faces and flowing through the sample (Sutherland & Cave, 1980). To understand the effect of porosity reduction on the permeability of young, loosely consolidated rocks, we measured FB12 permeability before and after mechanical compaction, which was assumed to be isotropic. First, we measured ultrasonic velocities and permeabilities at Pc up to 70 MPa and = 5 MPa, then, we increased Pc stepwise to 100, 150 and 200 MPa and waited for 19, 24 and 5 hours to measure creep until the observed shortening rate was less than 1 µm/hour. Finally, we measured the sample permeability for varying Pc up to 200 MPa.
To study how fractures influence the permeability of HM rocks, we split sample MT07 through a Brazilian test producing a sub-vertical fracture connecting the opposite end-faces of the core plug. Then, to study the effect of stress on permeability healing, we kept the sample dry and measured permeability as a function of s’ and we collected three micro-computed tomographies (mCT) to seek evidence of variations in fracture aperture. A detailed chronology of the operations follows: On day 1, after the Brazilian test, we collected mCT dataset S1. Between day 2 and 9 we performed the first permeability test (kT1) for s’ between 24 and 65 MPa. During kT1 (days 3 to 5) we promoted healing by keeping s’ to 65 MPa. After kT1 and for the next 39 days, the sample remained inside the pressure vessel at s’~0 MPa. Between day 48 and day 77, we performed the second permeability test (kT2) at s’ ranging 5.6 to 64 MPa. At the end of kT2 we removed the sample from the pressure vessel and acquired mCT dataset S2. Then, the jacketed sample was placed inside a humidity-controlled chamber equipped with a water container and a thermo-hygrometer. For 72 hours, a medium to low vacuum (<0.5 bar) was maintained to promote water evaporation, causing the chamber relative humidity to remain above 97% and activating clays such as smectites with pronounced swelling properties (Villar et al., 2005). Finally, we acquired mCT dataset S3, and produced a thin section perpendicular to the sample axis. On the thin section, we examined the morphology of the fracture for evidence of clay infilling, possibly caused by plastic deformation and triggered by clay swelling.
Each µCT dataset comprises 1600, 33.3 mm resolution, 16-bits TIFF images perpendicular to the sample axis, recording the entire sample except 4.37 mm at the top and bottom. After normalization and segmentation, we calculated fracture apertures (B) for each CT dataset by producing fracture aperture distribution projections (FADP) whose mean and standard deviation provided average apertures (Bm) and associated uncertainties. We report more details on the methods in the supporting information.
Figure 1. Geologic map of the Raukumara peninsula with the position of the rock samples used in this study (Mazengarb & Speden, 2000). The offshore dashed line contour marks the 50 mm geodetic slip model for the September-November 2014 SSE (Warren-Smith et al., 2019). The offshore line indicates the seismic line MC10 from the SHIRE project (Gase et al., 2021).
3 Results
The four samples (Fig S1) contain more than 35 wt% quartz and feldspars. The remaining minerals are calcite, and clays: chlorite, kaolinite, micas, illite, and smectite group minerals (Fig S2). Clays and swelling clays (i.e., illites and smectites) represent at least 24 wt% and 13 wt%, respectively (Fig S3). Porosities vary between 7 and 18%, where the tighter samples (MT07 and MO02) have a longer diagenetic or metamorphic history. Microphotography reveals that the grain size varies significantly among the four samples: Sample GB13 has the smallest grain size (<20 mm).
Ultrasonic velocity measurements (Fig 2A) show that Vp and Vs increase with sm, and the younger samples (FB12 and GB13) generally have lower wave speeds. Vp to Vs ratios vary between 1.6 and 1.95, with the least consolidated and youngest sample (FB12) exhibiting the highest values. After saturation, sample GB13 Vp increased by ~250 m/s on average while Vs decreased by ~100 m/s on average, increasing the Vp to Vs ratio from ~1.65 to ~1.95.
Figure 2. A) Left-top panel: ultrasonic velocities measured on the samples MO02, MT07 and FB12 as a function of sM. Right-top panel: ultrasonic Vp and Vs for sample GB13 when dry and saturated with water. Dashed lines indicate the theoretical saturated velocities from Gassmann fluid substitution (Gassmann, 1951). The inset shows examples of P and S waveforms recorded for the dry sample at the conditions indicated by the circled dots. Bottom panels: Vp to Vs ratios for the laboratory data and the fitting curves reported in the panels above. B) Permeabilities for samples FB12, MT07, and MO02 as a function of s'. (a) “FB12 compaction” reports the loss of permeability due to the step-by-step increase of s'; partial compaction and loss of porosity in sample FB12 are shown in "FB12 creep" and “FB12 compaction” insets, respectively. "Pp equilibration" inset: example of pore pressure (Pp) equilibration and fitting curves (dashed lines). The blue and red curves show the permeability of the fractured sample MT07 during the two permability cycles (kT1 and kT2). (d X) near data points indicates X days since stage S1 (Fig 3A). Day 77 was the end of kT2 and stage S2 (Fig 3B): the sample was CT-scanned and exposed to a humid atmosphere for 72 hours. Day 81 was stage S3: we CT-scanned and remeasured the sample permeability (Fig 3C,D,E).
Before compaction, sample FB12 permeability ranged between 200 and 400 mD. Then, we raised Pc twice to 70 MPa, causing the permeability to decrease by a factor of two and porosity by 3% (i.e., at sm~7 MPa, porosity varied from 17.3 to 14.2%). In the following two cycles, where Pc reached 200 MPa, porosity decreased to 13.9%, and the permeability declined by almost an order of magnitude. Concurrently, the ultrasonic Vp increased from 2.6 km/s to 4 km/s.
Samples MO02 and MT07, when intact, have permeabilities below 100 nD, regardless of s'. The permeability of the fractured MT07 evolved between stages S1, S2, and S3. After S1 and during the permeability cycle kT1, the permeability dropped from 2 mD to 87 nD. After exposing the sample to s' ~65 MPa for more than 48 hours (Fig 4B b), we continued kT1 and found that the permeability further decreased to 24 nD. The permeability remained ~2 orders of magnitude lower than the initial permeability, i.e., around 30 nD, when s' was reduced. After 39 days, the new increase of s' during the second permeability cycle kT2, caused the permeability to drop to 9 nD. During the following decrease of s', the permeability resembled pre-fracturing values. The last measurement of kT2 was performed at s'=4.5 MPa and permeability was 300 nD, seven times lower than the initial value measured at s'=5.6 MPa. After exposing the sample to humidity for 72 hours, the permeability, measured at s'=3.7 MPa, decreased to 67 nD.
CT-scans visual inspection and analyses reveal the variation of Bm that varied from 190+/-110, to 72+/-43 and 84+/-45 µm during the stages S1, S2, and S3, respectively (Fig 3A,B,C). During the same stages, the number of voxels counted within the fracture varied respectively from ~934,000 to ~581,000 and ~797,000. Microphotography of sample MT07 at stage S3, shows that in several loci, the fracture collapsed, and a fine-grained amorphous mass infilled the fracture (Fig 3D,E). These observations suggest that varying confining pressure and humidification caused clay minerals plastic deformation and swelling, partially closing the fracture and reducing the permeability.
Figure 3. CT-scan and transmitted light microphotography of sample MT07 after fracturing. A, B, and C are CT-scans at stages S1, S2, and S3, respectively. Each top-left inset in these panels reports a section of the CT-scan model after normalization. CT-number distribution is shown in the bottom-left inset. The red vertical line indicates (eq. S5). The percentages on the left and right of the red line indicate the relative quantity of voxels representing air and solid rock, respectively. The right top inset in each panel shows the binarized 3D model, where voxels collected within the fracture are blue. The bottom right inset show the aperture distribution (B), the calculated average and standard deviation (Bm and horizontal red bar), and the total count of voxel within the fracture (n). Panel D is a microphotography of the thin section at stage S3. Panel E reports zooms from panel D. Insets 1 to 4 show fracture infill, which are highlighted by red arrows along with open fractures.
4 Discussion
We provide porosity-permeability relationships for rock samples from the subaerial northern HM under a range of confining pressures. Ultrasonic velocities of dry samples are similar to the seismic velocities estimated offshore New Zealand by the SHIRE project (Gase et al., 2021). The seismic reflectivity imaged along the transect MC10 shows the decollement along the prism base and several splay faults that may partly accommodate the convergence (Fig 4A). Inside the prism, Vp increases gradually from 2.0 km/s near the surface to 4.5 km/s at the prism base ~7 km below sea level. In Fig 4B, the comparison between the seismic and ultrasonic velocities suggests that sample FB12, and possibly also sample GB13, represent the modern slope basins on the outer prism, which is consistent with their depositional environment. The ultrasonic velocities of sample MT07 of the Tinui Group correspond well to the velocities of the deep part of the prism, where Vp reaches 4.5 km/s. Compaction and diagenesis must contribute to the increase of Vp with depth (Dvorkin & Nur, 1996; Saxena & Mavko, 2014). We measured an ultrasonic Vp of 4.8 km/s at 150 MPa in the Torlesse basement sample MO02, which is higher than what we imaged in the deep prism on Line MC10 (Fig 4), suggesting that there may not be a deep offshore portion of the Torlesse basement offshore northern HM (Bassett et al., 2022; Gase et al., 2021).
Our comparison between seismic and ultrasonic velocities in Figure 4 is semi-quantitative as uncertainty is introduced by microcracks produced during sample preparation – see SI for details (Eberhart‐Phillips et al., 1989; Tsuji & Iturrino, 2008), and by frequency differences. Velocities in section MC10 and on our samples have been measured at frequencies around 20 Hz and 800 kHz, respectively. Considering the frequency range, a typical P-wave quality factor ranging from 30 to 150, and a nearly-constant Q model (Liu et al., 1976; Tisato et al., 2021), we should expect a velocity dispersion between 2.3 and 12%. Conversely, SHIRE and laboratory data were collected on saturated and dry samples, respectively. Saturation increases P-wave velocities of sample GB13 by 5 to 15%, suggesting that the effects of fluid saturation and anelasticity on velocities should counteract each other. Given the similarity in P-wave velocities and depositional environment, we suggest that the Tinui and Tolaga group rocks (samples MT07 and GB13) are good lithological proxies for the deep and shallow offshore Hikurangi prism, respectively.
Figure 4. A) Velocity model along the SHIRE Line MC10 (Gase et al., 2021). B) Summary of laboratory result: permeabilities vs porosity and color-coded markers (colorbar in panel A) as a function of ultrasonic Vp for samples FB12, MO02, and MT07 (Tables S1 and S2). The arrow indicates in which direction the permeabilities vary when tests are performed using water rather than helium gas. Dashed lines indicate empirical permeability vs porosity according to eq. 3. The dotted line represents an average permeability for unconsolidated clays and possibly a lower bound for the permeability of HM sediments (Neuzil, 1994). S data (dark-gray area) are for siltstones (Reece et al., 2012). The continuous line fits our data and agrees with measured mudstone permeabilities indicated by the MN gray-shaded area (Magara, 1978; Neglia, 1979). Such a line also represents an upper bound for the permeability of HM rocks. M, T, and B data are permeabilities measured in boreholes: M by Reisdorf et al. (2016), Yu et al. (2017); T by Boisson et al. (2001); B by Intera Eng. Ltd. (2011), Roberts et al. (2011), Walsh (2011).
The permeability of our samples ranges from 1 nD to 1 mD, with the samples representing the deep part of the prism being the tightest. Neuzil (1994, 2019) compiled data from several studies on unconsolidated clays with a maximum porosity of 80%, and a few consolidated mudstone-siltstones with porosities (F) <35%. Saffer & Bekins (1998) followed Neuzil’s work and described the permeability (k) of the Nankai accretionary complex as:
eq. 1
Equation 1 fits the porosity-permeability relationship of unconsolidated sediments and is a lower bound for the permeability of mudstones that are similar to our samples (Magara, 1978; Neglia, 1979; Reece et al., 2012). On the other hand, we found that:
eq. 2
fits our results and is an upper bound for the permeability of mudstones. We suggest that the permeabilities calculated from equations 1 and 2 (Fig 4B) overestimate permeabilities in the Northern Hikurangi accretionary prism at depths >1 km because helium gas is not as efficient as seawater in activating swelling clays whose expansion lowers the effective permeabilities (Villar et al., 2005); At burial depths >1-2 km, the porosity of clay-bearing sediments, mudstones, siltstone, and shales drops below 35% (Griffiths & Joshi, 1989; Magara, 1978; Skempton, 1969); Permeabilities measured in boreholes are typically orders of magnitude higher than those measured in the laboratory due to the presence of fractures (Fig 4B lines M,T,B) (Neuzil, 2019), and numerical models of permeability in microfractured claystones agree with the mudstone porosity-permeability in Fig 4B (Vora & Dugan, 2019). We also propose that the permeability of rocks in the Northern Hikurangi accretionary prism can be described by a Kozeny-Carman relation (dashed lines in Fig 4B):
eq. 3
Where is tortuosity, and R is the median pore diameter (Carman, 1997). We obtained from data reported by Hunt (1996) for similar lithologies.
Every 1-2 years, the northern HM experiences an SSE that lasts several weeks (Wallace, 2020). Recent analyses of the APG data offshore Gisborne have shown that the 2014 SSE may have experienced up to 30 cm of slip in the center of a ~100 km wide patch, though less displacement is expected along the edges (Yohler et al., 2019). Some authors have suggested that SSEs that originate along the decollement at the base of the wedge are accompanied by slip diverted to thrust faults in the Hikurangi accretionary wedge (Shaddox & Schwartz, 2019). We expect SSEs to deform and fracture the rocks along these thrust faults (Morgan et al., 2022). Our laboratory measurements before and after rock failure for sample MT07 show that the deeper prism, where Tinui Group equivalent rocks may be present, may experience large increases in permeability during an SSE.
In a few weeks, the fractured sample MT07 regained its pre-fracturing permeability. Between stages S1 and S2, the permeability recovery was achieved in dry conditions. Although sample MT07 and sample FB12 have different compaction levels and grain sizes, they share similar mineralogy. Thus, although limited, we expect plastic deformation also in sample MT07, likely concentrated near clays (Mondol et al., 2008). Between stages S2 and S3, the permeability decreased by a factor of 5 while Bm increased, suggesting clay expansion. Once confined, we expect that the hydrated clays would deform plastically, clogging the fracture more efficiently than dry clays and justifying the permeability loss. We propose that permeability healing is also present along HM faults, given the presence of clays at depth, especially above the 5-7 km deep temperature-controlled smectite-illite transition (Antriasian et al., 2018; Freed & Peacor, 1989; Pecher et al., 2017; Tisato & Marelli, 2013).
In the Hikurangi subduction zone, fluids expelled from pore space and fluids released by dehydration reactions travel along the plate interface or through the accretionary wedge (Ellis et al., 2015). As the fluid pressure increases near the decollement and inside the accretionary wedge, conditions may become favorable for an SSE (Burgreen-Chan et al., 2016; Kobayashi & Sato, 2021). Though this mechanism has been proposed for several subduction zones where SSE occur at larger depths (Audet et al., 2009; Kodaira et al., 2004), the analysis of Warren-Smith et al. (2019) on the northern HM, is also compatible with the sealing of fluid pathways after an SSE. The expansion and plastic deformation of clays may provide an efficient mechanism to reduce permeability over weeks or months after an SSE.
Permeability healing, favoring the development of overpressures, reconciles with the poor mechanical healing shown by Shreedharan et al. (2023), hindering elastic energy accumulation, because both set conditions conducive to SSEs.
5 Conclusions
We provided relationships between porosity, permeability, and confining pressure for rocks that make up the accretionary prism of the northern HM. We suggest an empirical porosity-permeability relationship to model fluid transport and estimate effective stress in shallow subduction zones. Mechanical failure of these rocks enhances permeability, but over the course of several weeks, healing reduces the permeability again, suggesting that after an SSE, sediments deep in the northern HM accretionary prism can recover permeability efficiently within the time frame of an SSE as a mechanism explaining the regular recurrence of these events.
Acknowledgments
We thank Lucy Flesh, Carolyn Boulton, and one anonymous reviewer for the insightful reviews, Luc Lavier and Dominic Strogen for the valuable discussions. NSF supported this research (grant OCE-1949171).
Open Research
Data are publicly available upon publication at https://doi.org/10.18738/T8/RMXMIQ or can be requested to the corresponding author.
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