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Faulting Processes Unveiled by Magnetic Properties
of Fault Rocks
Tao Yang
1,2,3
,Yu‐Min Chou
4,5
, Eric C. Ferré
6
, Mark J. Dekkers
2
, Jianye Chen
7,8
,
En‐Chao Yeh
9
, and Wataru Tanikawa
10
1
Hubei Subsurface Multi‐Scale Imaging Key Laboratory, Institute of Geophysics and Geomatics, China University of
Geosciences, Wuhan, China,
2
Paleomagnetic Laboratory Fort Hoofddijk, Department of Earth Sciences, Utrecht
University, Utrecht, The Netherlands,
3
Now at School of Geophysics and Information Technology, China University of
Geosciences, Beijing, China,
4
Shenzhen Key Laboratory of Marine Archaea Geo‐Omics, Department of Ocean Science and
Engineering, South University of Science and Technology, Shenzhen, China,
5
Southern Marine Science and Engineering
Guangdong Laboratory, Department of Ocean Science and Engineering, Southern University of Science and Technology,
Shenzhen, China,
6
School of Geosciences, University of Louisiana at Lafayette, Lafayette, LA, USA,
7
State Key Laboratory
of Earthquake Dynamics, Institute of Geology, China Earthquake Administration, Beijing, China,
8
HPT Laboratory,
Department of Earth Sciences, Utrecht University, Utrecht, Netherlands,
9
Department of Earth Sciences, National Taiwan
Normal University, Taipei, Taiwan,
10
Japan Agency for Marine‐Earth Science and Technology, Kochi Institute for Core
Sample Research, Nankoku, Japan
Abstract As iron‐bearing minerals—ferrimagnetic minerals in particular—are sensitive to stress,
temperature, and presence of fluids in fault zones, their magnetic properties provide valuable insights into
physical and chemical processes affecting fault rocks. Here, we review the advances made in magnetic
studies of fault rocks in the past three decades. We provide a synthesis of the mechanisms that account for
the magnetic changes in fault rocks and insights gained from magnetic research. We also integrate
nonmagnetic approaches in the evaluation of the magnetic properties of fault rocks. Magnetic analysis
unveils microscopic processes operating in the fault zones such as frictional heating, energy dissipation, and
fluid percolation that are otherwise difficult to constrain. This makes magnetic properties suited as a
“strain indicator,”a“geothermometer,”and a “fluid tracer”in fault zones. However, a full understanding of
faulting‐induced magnetic changes has not been accomplished yet. Future research should focus on
detailed magnetic property analysis of fault zones including magnetic microscanning and magnetic fabric
analysis. To calibrate the observations on natural fault zones, laboratory experiments should be carried
out that enable to extract the exact physicochemical conditions that led to a certain magnetic signature.
Potential avenues could include (1) magnetic investigations on natural and synthetic fault rocks after
friction experiments, (2) laboratory simulation of fault fluid percolation, (3) paleomagnetic analysis of
postkinematic remanence components associated with faulting processes, and (4) synergy of
interdisciplinary approaches in mineral‐magnetic studies. This would help to place our understanding of the
microphysics of faulting on a much stronger footing.
Plain Language Summary The Earth's surface is riddled with faults that largely contribute to
landscape evolution and human activities. Some of these faults produce earthquakes of different magnitudes
including some with catastrophic consequences. Understanding faulting mechanisms benefits society
when predictions about rupture are made. Fault zones preserve an excellent record of chemical and physical
processes involved in failure. Among other analytical methods, magnetic studies prove to be an
emergent and untapped source of information on these processes. These methods, focused on prefaulting,
synfaulting, and postfaulting mineral changes, have resulted in significant advances in our
understanding of the conditions of faulting. In this review, we present an extensive account of the state of
knowledge and highlight current challenges and future avenues of fault magnetism research.
1. Introduction
Fault zones comprise only a very small volume of the Earth's crust. However, their structure and associated
deformation processes are ultimately decisive for a wide range of crustal processes (e.g., Faulkner
et al., 2010; Townend & Zoback, 2000). This pertains in particular to the physical origin of earthquakes
©2020. The Authors.
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terms of the Creative Commons
Attribution License, which permits use,
distribution and reproduction in any
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properly cited.
REVIEW ARTICLE
10.1029/2019RG000690
Key Points:
•Iron‐bearing minerals and
ferrimagnetic minerals in particular
are sensitive to faulting‐associated
physical and chemical processes
•Laboratory faulting experiments and
comparison with nonmagnetic
approaches confirm results from
magnetic studies on natural rocks
•Rock magnetic methods offer
novel tools to analyze strain, grain
fining, temperature trends, and
fluid‐rock interaction in fault zones
Correspondence to:
T. Yang and M. J. Dekkers,
tyang@cugb.edu.cn;
t.yang@uu.nl;
m.j.dekkers@uu.nl
Citation:
Yang, T., Chou, Y.‐M., Ferré, E. C.,
Dekkers, M. J., Chen, J., Yeh, E.‐C.,
& Tanikawa, W. (2020). Faulting
processes unveiled by magnetic
properties of fault rocks. Reviews of
Geophysics,58, e2019RG000690.
https://doi.org/10.1029/2019RG000690
Received 28 FEB 2020
Accepted 21 SEP 2020
Accepted article online 3 OCT 2020
YANG ET AL. 1of60
within the seismogenic zone, which typically extends from 3–4to15–20 km depth for faults in the con-
tinental crust (e.g., Scholz, 2019; Sibson, 1986). “Fault‐related rocks”(Wise et al., 1984), or simply “fault
rocks”(Sibson, 1977), form as a result of localized strain within a fault zone (Brodie et al., 2007). These
rocks are common along upper crustal fault zones (Woodcock & Mort, 2008). Physical and chemical
attributes, and textures of fault rocks hold valuable information for understanding both the long‐term
behavior of faults, on a timescale of millions of years, as well as the short‐term behavior involving the
nucleation (see Appendix A for definition), rupture, cessation, and recurrence of (large) earthquakes
(Bradbury et al., 2015; Faulkner et al., 2010; Henderson et al., 2010; Rowe & Griffith, 2015; Schmid &
Handy, 1991; Ujiie & Kimura, 2014; Wibberley et al., 2008). Studies into fault rocks thus have
intrigued geologists for decades and have been and continue to be an important subject in the context
of large, active faults with repeated seismic activity (Chester et al., 1993; Rowe & Griffith, 2015;
Sibson, 1986).
It is realized that a full appreciation of earthquake dynamics requires an integration of macroscopic seismol-
ogy with microscopic and experimental studies of fault zones (e.g., Cowan, 1999; Kanamori & Heaton, 2000;
K.‐F. Ma, 2009; Niemeijer et al., 2012; Scholz, 2019; Sibson, 1989). The former relies on earthquake kine-
matic analysis based on seismic waveform data (Kanamori & Heaton, 2000). The latter two mainly focus
on the physical/chemical attributes and texture of fault rocks, including their formation mechanisms
(K.‐F. Ma, 2009; Scholz, 2019). Their combination is crucial for understanding earthquake energy dissipa-
tion, rupture processes, and seismic efficiency. Substantial research effort has been dedicated during the
last two decades to understanding the development and physical/chemical attributes of fault rocks using
diverse approaches—either of mesoscopic or microscopic scale—in natural faults, laboratory experiments,
and numerical simulations (e.g., Billi, 2005; Mair & Abe, 2008; Rowe, Fagereng, et al., 2012; Rowe,
Kirkpatrick, & Brodsky, 2012). These studies have covered aspects of fault rocks in the widest possible sense,
including terminology and classification (Choi et al., 2016; Woodcock & Mort, 2008), development and
architecture (Faulkner et al., 2010; Fossen & Rotevatn, 2016; Pei et al., 2015; Wibberley et al., 2008), micro-
structure (e.g., Boullier et al., 2009; Bradbury et al., 2015; Fossen & Cavalcante, 2017; Isaacs et al., 2007;
Rowe & Griffith, 2015), grain‐size distribution (e.g., Billi, 2005; Hattori & Yamamoto, 1999; Wilson
et al., 2005), mineral assemblages (e.g., Boullier, 2011; Bradbury et al., 2015; J. Chen et al., 2013; Isaacs
et al., 2007; Matsuda et al., 2004), geochemical composition (e.g., Bradbury et al., 2015; J. Chen et al., 2013;
Isaacs et al., 2007; Tanaka et al., 2001), hydraulic properties (e.g., Bense et al., 2013; Carpenter et al., 2014;
J. Chen et al., 2016; Faulkner et al., 2010), nanocrystallization (e.g., Verberne et al., 2019; Viti, 2011),
geochronology (e.g., Oriolo et al., 2018), and frictional properties (e.g., Boulton et al., 2017; Carpenter
et al., 2015; Di Toro et al., 2011; Niemeijer & Vissers, 2014).
A magnetic property analysis is occasionally included in the description of fault rock properties. Magnetic
properties are primarily governed by the distribution and speciation of iron—the fourth most abundant ele-
ment in the Earth's crust. By far most Fe‐bearing minerals are not considered “magnetic”at room tempera-
ture; common minerals include Fe‐bearing silicates, carbonates (siderite and ankerite), and sulfides (pyrite).
What we refer to as magnetic minerals are ferromagnetic (sensu lato) minerals, which retain permanent or
remanent magnetism; they include iron oxides (i.e., (titano)magnetite, maghemite, and hematite), oxyhydr-
oxides (primarily goethite), and some iron sulfides (i.e., pyrrhotite and greigite). They occur only in trace
amounts in the rock but nonetheless determine its magnetic signature. Metallic iron and other reduced
phases such as iron phosphide, iron carbide, and iron silicide are also magnetic; they are extremely rare
in the Earth's crust. All aforementioned minerals, magnetic and nonmagnetic, are not exempt from
(thermo)chemical alterations induced by the faulting and related fluid‐rock interaction, resulting in changes
in magnetic properties of fault rocks in comparison to adjacent host rocks. Thus, magnetic properties of fault
zones are natural archives of the faulting‐associated processes in tectonically active regions. It makes rock
magnetism, which is the study of the magnetic properties of rocks, sediments, soils, and even organisms,
a promising tool to unravel faulting processes. Rock magnetic techniques are capable of determining the nat-
ure, grain size, and concentration of magnetic minerals in a sample down to the ppm level, which lies well
below the detection limit of more conventional mineralogical techniques, for example, X‐ray diffraction.
Also, magnetic measurements are generally rapid, cost‐effective, and nondestructive (evidently with the
exception of high‐temperature magnetic analysis) so they can be used in conjunction with other techniques.
These advantages make rock magnetic analysis of fault rocks attractive.
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Fault rocks, however, have been the subject of rock magnetic studies only recently, with a focus on faults
from seismically active zones (e.g., Almqvist et al., 2020; Cai et al., 2019; Chou, Song, Aubourg, Lee,
et al., 2012; Chou, Song, Aubourg, et al., 2014; Chou, Song, Aubourg, Song, et al., 2012; Chou, Song, Lee,
et al., 2014; Ferré et al., 2005, 2012, 2015, 2016, 2017; Ferré, Gébelin, et al., 2014; Ferré, Geissman, et al., 2014;
Fukuchi, 2003; Fukuchi et al., 2005, 2007, 2009; D. Liu et al., 2014, 2016; Mishima et al., 2006,
2009; Nakamura et al., 2002; Pei, Li, et al., 2014; Pei, Zhou, et al., 2014; Tanikawa et al., 2007, 2008; Yang
et al., 2012a, 2012b, 2018, 2019; Yang, Chen, et al., 2013; Yang, Mishima, et al., 2013; Yang, Dekkers, &
Zhang, 2016; Yang, Yang, et al., 2016; Zhang et al., 2017, 2018). By providing a trove of information on phy-
sical and chemical processes associated with faulting in seismically active zones, these pioneer works have
laid the foundation of “fault magnetism”, an emerging subdiscipline in the geosciences.
This review concentrates on the rock magnetic studies on fault rocks from seismically active fault zones
where recent megaearthquakes occurred, such as the 1995 Mw (moment magnitude) 6.9 Kobe (Japan),
the 1999 Mw 7.6 Chi‐Chi (Taiwan), the 2008 Mw 7.9 Wenchuan (China), and the 2011 Mw 9.0 Tohoku
(Japan) earthquakes. Additionally, magnetic studies on pseudotachylytes that are the products of fossil seis-
mogenic fault zones are included. First, we describe concisely the processes that result in the formation of
fault rocks. Second, we outline the approaches to evaluate magnetism of fault rocks. Third, we discuss the
mechanisms that are responsible for changes in the magnetism of fault rocks. Fourth, we address the
insights gained from magnetic property analysis of fault rocks and compare the magnetic methods with
other approaches, the core of this review. Thereafter, we identify some of the current challenges in fault mag-
netism research and propose potential avenues that would contribute to a full appreciation of the magnetism
of fault rocks and would advance our understanding of faulting.
2. Formation of Fault Rocks and Processes During Faulting
A fault is a narrow zone of crushed rocks along which two blocks of rock have moved alongside each other in
response to stress imposed on the rock. This movement, that is, the very faulting, takes place either as creep:
stable, slow sliding, or in the form of a series of earthquakes: unstable, fast slip intertwined with long periods
of no motion. Faults thus accommodate strain on a momentous range of dimensions: on a spatial scale, from
millimeters to hundreds of kilometers (for major plate boundaries), while also the time intervals range
widely from seconds to minutes of earthquake slip, to years of slow, aseismic slip, and to millions of years
of intermittent activity (e.g., Nielsen, 2017; Sibson, 2003). The textures and structures that develop in fault
rocks depend on the amount and rate of shearing, and the physical conditions, that is, temperature and pres-
sure, under which the shearing occurred. Fault zones that develop within and above the seismogenic zone
are often characterized by highly localized brittle shear deformation (Figure 1a); they are termed brittle
faults. In contrast, faults formed at much deeper levels in the crust tend to deform by purely thermally acti-
vated creep mechanisms and are characterized by continuous ductile displacement; they are termed ductile
faults (e.g., Scholz, 2019). As depth increases, a fault zone is expected to encounter a gradual transition from
brittle to ductile deformation, occurring in a depth interval with temperatures spanning from a few tens to
hundreds of °C (typically 400–650°C, depending on the fault rock composition). It is usually termed the
“brittle‐to‐ductile”transition zone (e.g., Aharonov & Scholz, 2019; Kawamoto & Shimamoto, 1997). The
“brittle‐to‐ductile”transition zone also defines the lower limit of seismic activity on faults since ductile
deformation is largely aseismic (Aharonov & Scholz, 2019).
The terminology and classification of fault rocks are not universally agreed upon yet (Woodcock &
Mort, 2008). Fault rocks formed above a depth of 1–4 km are generally incoherent, friable, and uncemented;
they include fault breccia and fault gouge (Figure 1a). At greater depth (up to 10–15 km), typical fault rocks
are different types of cataclasites, sometimes with pockets of pseudotachylyte that were formed during seis-
mic slip events (e.g., Sibson & Toy, 2006). The cataclasite group is subdivided according to the relative pro-
portion of fine‐grained matrix into protocataclasite, (meso)cataclasite, and ultracataclasite (Figure 1b; e.g.,
Brodie et al., 2007; Woodcock & Mort, 2008). A similar classification also applies to mylonite as the fault rock
type for ductile deformation at greater depth (Figure 1b). The “brittle‐to‐ductile”transition zone, as the
name suggests, is characterized by mixed mode of fault slip behavior and by extensive juxtaposing and over-
printing of different types of fault rocks, such as mylonite with pseudotachylyte interlayers and mylonitized
pseudotachylyte (e.g., Hayman & Lavier, 2014; Menegon et al., 2017; White, 1996).
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Figure 2 summarizes the formation stages of a typical brittle fault (e.g., Hattori & Yamamoto, 1999):
(i) increased stress produces a cross‐joint system over the rock body (Figure 2a); (ii) rock fragments are
formed by shearing along joints (Figure 2b); (iii) movements of the host blocks in a bookshelf‐style mode
rotate these fragments and abrase them to sand‐, silt‐, and clay‐size particles, even down to micron/nan-
ometer‐size particles (Figure 2c); (iv) cementation of disintegrated fragments and subsequent formation of
new joints (Figure 2d) with one or more cycles of refracturing (Figure 2e); and eventually (v) a relatively nar-
row zone (millimeters to centimeters thick) composed of very fine grained particles is produced (Figure 2f).
The fault growth process commonly results in a fault zone architecture consisting of a so‐called fault core
enclosed by broader damage zones (Figure 2g; Caine et al., 1996; Chester et al., 1993).
The fault core is the result of highly localized strain and intense shearing that accommodates the majority of
the displacement within the fault zone. It typically consists of a number of (recurring) slip surfaces and sev-
eral types of fault rocks, such as fault breccia, fault gouge, and/or cataclasite (Caine et al., 1996; Chester
et al., 1993; Choi et al., 2016; Sibson, 2003). The principal slip zone (PSZ) cuts the fault core (Figure 2g); it
is the foremost location where physicochemical processes take place as a result of individual earthquake
events (Boullier, 2011; Sibson, 2003). Examples of fault rocks from the first borehole of the Wenchuan earth-
quake Fault Scientific Drilling project (WFSD‐1) are shown in Figure 2 (Wang et al., 2014).
The damage zone enveloping the fault core (Figure 2g) is characterized by what is referred to as distributed
deformation (i.e., where homogenous strain is distributed across a network of evenly spaced faults, cf. Nixon
et al., 2014). The damage zone differs structurally, mechanically, and petrophysically from the undeformed
host rock or protolith (Caine et al., 1996; Chester et al., 1993; Choi et al., 2016). The host rock surrounds the
damage zone and remains basically unchanged during faulting (Caine et al., 1996). As mentioned earlier on,
a fault undergoes cycles of creep and seismic slip. A seismic cycle is divided into three periods: (1) the coseis-
mic period (the time during an earthquake itself, typically seconds to minutes), (2) the postseismic period
(days, months, and sometimes years after a given earthquake), and (3) the interseismic period (the time
Figure 1. (a) Sketch illustrating the occurrence of diverse fault rocks in the shallow crust (reproduced from
A. Lin et al., 2005). Fault breccia, fault gouge, and cataclasite mostly occur in the brittle regime; in contrast, mylonites
develop within the ductile regime. Pseudotachylytes are frozen melt pockets occasionally formed by seismic frictional
heating during large earthquakes. (b) A revised classification of fault rocks as suggested by Woodcock and Mort (2008).
The brittle versus ductile behavior of rocks depends, to a large extent, on strain rate. The delineation of brittle and
ductile zones here reflects expected behavior under typical geological strain rates of 10
−14
s
−1
.
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Figure 2. (a–f) An idealized schematic model showing the development of a brittle fault zone (modified after Hausegger et al., 2010, and Billi, 2005) and
(g) the typical architecture of a fault zone (after Chester et al., 1993, and Caine et al., 2010). A brittle fault zone is the result of several deformation processes:
(a) formation of cross joints, (b) fracturing of initial rocks and fragmentation, (c) fracturing and disintegration of fragments by bookshelf rotation, (d) cementation
of disintegrated fragments, (e) subsequent formation of new joints due to the refracturing, and (f) eventually a relatively narrow zone (i.e., principal slip
zone, PSZ) is formed composed of very fine grained particles. (g) Typical fault zone structure consisting of the fault core that includes fault gouge cut by the PSZ,
fault breccia, and/or cataclasite, the damage zone surrounded by undeformed host rock (protolith). The figure is not to scale. (h) Examples of different fault
rocks, which were retrieved from the first borehole of the Wenchuan earthquake Fault Scientific Drilling project (WFSD‐1) (adapted from Wang et al., 2014). The
blue and red lines indicate the boundary between cataclasite and fault gouge, and the inferred PSZ of the 2008 Mw 7.9 Wenchuan earthquake, respectively.
The numbers on both sides of the cores indicate the depth intervals they were retrieved from.
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span between large earthquakes, lasting tens to thousands of years). A fault zone is the cumulative product
of multiple slip cycles (Figure 3). In the following, we summarize some of the processes active in fault zones
during the seismic cycle with bearing on the magnetic properties of fault zones.
Thermochemical reactions and pseudotachylytes. During an earthquake (Figure 3a), the mechanical energy
due to the high seismic slip rates is converted into heat; it constitutes the largest part (~80% to 90%) of the total
earthquake energy budget (Pittarello et al., 2008; Scholz, 2019). Frictional heating will quickly raise the tem-
perature in the slip zone after rupture (Rice, 2006). Considering typical seismic slip rates (1 m/s) and total slip
distance (tens of centimeters to meters) at a fault plane, the related temperature rise is <100°C near Earth's
surface but can become >1100°C at seismogenic depths (>5 km) (e.g., McKenzie & Brune, 1972). The tem-
perature rise may promote thermal decomposition or dehydration of certain mineral phases and the forma-
tion of breakdown products (e.g., Di Toro et al., 2011, Figure 3a). This also pertains to the magnetic minerals
(see section 4.3 for a detailed discussion). The temperature rise within the slip zone is sometimes sufficient to
trigger melting of the host rock minerals (typically >1000°C, e.g., Spray, 2010). Quick, quench cooling of
these local melt pockets of millimeter to decimeter scale produces glasses or partly devitrified glasses, called
pseudotachylytes (Figure 1a; e.g., A. Lin, 2008; A. Lin et al., 2005; McKenzie & Brune, 1972; Philpotts, 1964;
Figure 3. A conceptual model showing the faulting‐related physical and chemical processes and the causes of potential magnetic changes in a fault zone during
the different stages of the earthquake cycle. (a) During a coseismic period, the high rate seismic slip may crush the ferrimagnetic grains (if present) in the
PSZ (indicated by the dark band) to a much finer size, due to the comminution of rock fragments. Meanwhile, frictional heating raises the temperature and can
induce a myriad of thermochemical reactions within a fault zone, which include transformations of Fe‐bearing minerals with immediate bearing on magnetic
properties. Coseismic hot fluids and those from deeper down may dissolve the Fe‐bearing minerals and release mobile Fe
2+
. (b) During a postseismic period,
neoformation of ferrimagnetic minerals occurs through precipitation of the mobilized Fe
2+
as mixed ferrous‐ferric iron oxide. Further, several thermochemical
reactions may proceed during cooling of the fault fluids warmed by friction in the PSZ. (c) During an interseismic period, infiltration and percolation of
meteoric water and/or deep‐layer fluids into the fault zone may dissolve (part of) the Fe‐bearing minerals again and induce precipitation of new magnetic
minerals. A fault zone is the cumulative product of many earthquake cycles. The fault core cartoons are modified after Gray et al. (2005).
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Sibson & Toy, 2006; Spray, 2010; Swanson, 1992). One should realize that the wording pseudotachylyte is also
in use to describe any aphanitic, dark colored fault rock or amorphous material formed during shear defor-
mation (see Rowe & Griffith, 2015). Throughout this review, however, we restrict pseudotachylytes to rocks
with a frictional melt origin. Pseudotachylytes commonly have a relatively high concentration of fine‐grained
magnetite as a result of oxidation of melt‐susceptible mafic minerals (e.g., Ferré et al., 2005; Nakamura
et al., 2002; O'Hara, 2001; Pittarello et al., 2012; Zhang et al., 2018). This makes pseudotachylytes stand out
magnetically, in comparison with their host rocks, so that they are attractive rocks for magnetic characteriza-
tion (see section 4.4 for a detailed discussion).
Fluid movement in fault zones. Fault zones feature a dense network of fractures and secondary faults and thus
act as major fluid conduits in the crust (Bense et al., 2013; Vermilye & Scholz, 1998). Fluids of multiple
sources, including meteoric waters, trapped formation brines, mineral dehydration, and volatiles from the
deep underlying layers (Hickman et al., 1995; Zoback et al., 2007), can infiltrate into and percolate along fault
zones. Overpressurization may occur promoting earthquake nucleation (e.g., Miller, 2013; Scuderi &
Collettini, 2016; Sibson, 1992); pore fluid pressurization during frictional heating flashes may induce fault
weakening lengthening fault slip time spans (e.g., J. Chen et al., 2017; Miller, 2013; Rice, 2006).
Importantly, fluid‐related dissolution‐precipitation processes are anticipated to be common during all peri-
ods of the seismic cycle, that is, the coseismic, postseismic, and interseismic periods (Figure 3; e.g., Bense
et al., 2013; J. Chen et al., 2013, 2016; Gratier et al., 2013). This is discussed in section 4.5. These reactions play
a critical role not only in physical, chemical, and mechanical evolution of fault rocks (Bradbury et al., 2015;
J. P. Evans & Chester, 1995; Goddard & Evans, 1995; Isaacs et al., 2007; Niwa et al., 2015; Sutherland
et al., 2012) but also in the very earthquake rupture (Hickman et al., 1995; Sutherland et al., 2012; Zoback
et al., 2007).
Earthquake lightning. It is also reported that during large earthquakes (usually with Mw > 5), the crus-
tal deformation activates and releases gases and/or electrical charges. These subsequently generate
atmospheric electric fields and currents, which in turn affect the atmospheric electric circuit possibly
resulting in luminous phenomena (e.g., Derr, 1973; Enomoto et al., 2017; C.‐L. Kuo et al., 2011;
Lockner et al., 1983; St‐Laurent et al., 2006). This transient coseismic electric phenomenon is referred
to as earthquake lightning (EQL), which travels along the fault plane and may produce anomalous mag-
netizations in fault rocks (e.g., fault gouges and pseudotachylytes; Enomoto & Zheng, 1998; Enomoto
et al., 2001; Ferré et al., 2005). The mechanisms and magnetic effects are discussed in section 4.6.
All in all, a fault zone should be considered an extremely dynamic system, characterized by specific physico-
chemical conditions during each of the faulting stages (Figure 3; e.g., Cerchiari et al., 2020; Wibberley
et al., 2008). For example, the redox state associated with faulting and accompanying fluid flow varies enor-
mously during seismic slip events (Ishikawa et al., 2008; Yamaguchi et al., 2011). Also, fault rocks are subject
to multiple deformation events, often under different physicochemical conditions. Consequently, a fault rock
is not simply a granulated product of its protolith but is among the most complex and heterogeneous geolo-
gical materials. It may be viewed as a low‐to medium‐grade metamorphic rock (J. P. Evans & Chester, 1995).
Iron features three main valence states, zerovalent, ferrous, and ferric iron (along with several more exotic
intermediate forms) and is therefore a sensitive probe of the physicochemical conditions within fault zones.
Iron cycling as a consequence of faulting would cause deposition, leaching, and/or chemical alteration of the
Fe
2+
and/or Fe
3+
in Fe‐bearing minerals in fault zones. Very reducing conditions could lead to metallic iron.
In principle, this leads to measurable changes in magnetic properties, making rock magnetism a suitable
probing technique to decipher faulting processes. Unraveling the magnetic properties of fault rocks thus pro-
vides important clues to understand the evolution of fault zones. Also, aspects of earthquake physics and
chemistry may be unveiled. In the following section (section 3) we outline the merits of the rock magnetic
or mineral magnetic methodology (with emphasis on fault rocks) before detailing the drivers of magnetic
changes in fault rocks (section 4).
3. Rock and Mineral Magnetism of Fault Rocks
“Magnetic minerals”or ferromagnetic minerals (sensu lato) refer to minerals that are magnetic at room tem-
perature: iron‐titanium oxides, some iron sulfides, and metallic iron; the latter, however, rarely occurs in the
Earth's crust. Coupled electron spins over large atomic distances in the crystal structure result in collective
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spin behavior. Each magnetic mineral has a magnetic ordering temperature above which it loses its collective
magnetism and becomes a paramagnet; on cooling through the ordering temperature the collective magnetic
behavior is restored. At low temperatures below ~10–20 K, however, many more minerals order magnetically
and low‐temperature magnetic instrumentation can evaluate also those minerals.
The magnetic assemblage, concentration of each magnetic mineral, and magnetic granulometry of fault
rocks reflect faulting processes. The rock's magnetic features are sensitive to both chemical and physical
changes occurring in rocks during the faulting. For example, variations in physical grain size are often
reflected in the magnetic granulometry, which is the (inferred) grain‐size distribution of magnetic particles
in a sample, usually expressed through the dominant magnetic domain structure (Figure 4a). A particle's
domain structure is the result of minimizing the overall particle energy of individual magnetic particles.
The domain structure or state depends on the size and shape of the magnetic particles, next to being tem-
perature dependent. Crystal defects and the internal stress distribution in the magnetic grains also have their
impact on the domain state (Özdemir & Dunlop, 1997), which may be particularly relevant for rocks and
minerals that have undergone faulting.
We now provide (approximate) size ranges for the domain state types distinguished for magnetite, by far the
most common magnetic mineral in nature. At room temperature, the smallest particles, with a diameter of
up to 25 nm are superparamagnetic (SP); they cannot retain a geologically stable natural remanent magne-
tization (NRM). Next in size, in the 30–80 nm range (for equant grains), particles are single domain (SD).
Larger particles, up to a few μm in size, are featuring noncollinear spin structures and possibly even a
few domains; they are termed classically pseudo single domain (PSD, Stacey & Banerjee, 1974) but probably
more correctly should be referred to as “vortex state”particles (Schabes & Bertram, 1988; see also Almeida
et al., 2016, who show examples of particles in a vortex state) as recently proposed by Roberts et al. (2017). SD
and PSD or vortex particles are paleomagnetically most stable. Large grains, >5–10 μm or so, contain many
domains and are termed multidomain (MD). The reader is referred to Appendix A for further explanation;
textbooks include Dunlop and Özdemir (1997), M. E. Evans and Heller (2003), Stacey and Banerjee (1974),
and Tauxe (2010). Next to size also particle shape exerts a critical control on domain state threshold sizes. For
example, rod‐like particles of 0.1 × 0.1 × 1 μm are SD, whereas equant particles of 1 μm can be MD
(Figure 4a).
Various types of magnetic measurements and their combinations are used to determine the magnetic
carrier(s), their concentration, and (magnetic) grain size. They are categorized along field‐, frequency‐,
and temperature‐dependent measurements. Measurement of various types of laboratory‐induced remanent
magnetizations, that is, anhysteretic remanent magnetization (ARM) and isothermal remanent magnetiza-
tion (IRM), also yields important information. The fundamentals and applications of rock and mineral mag-
netism are documented in several textbooks (e.g., Dunlop & Özdemir, 1997; M. E. Evans & Heller, 2003;
Maher & Thompson, 1999; Nagata, 1961; Stacey & Banerjee, 1974; Tauxe, 2010; Thompson &
Oldfield, 1986) and review papers (e.g., Dekkers, 1997; Hunt et al., 1995; Q. S. Liu et al., 2012; Peters &
Dekkers, 2003; Verosub & Roberts, 1995). Below we outline the more relevant approaches for the analysis
of fault rocks. In a number of cases also nonmagnetic methods are utilized in concert with magnetic methods
to constrain the latter's interpretation. Before describing those magnetic and nonmagnetic methods, how-
ever, some specifics on sample collection and preparation for magnetic studies on fault rocks need to be con-
sidered. Magnetic property analysis requires physical samples since magnetic changes due to faulting
processes may be reflected on the mm to μm scale. This implies drilling and coring rocks through fault zones;
borehole logging “only”enables tracking broader (magnetic) aspects of fault zones.
3.1. Samples for Magnetic Studies Into Fault Rocks
As mentioned above, fault zones are narrow zones with widths in the order of centimeters to meters.
Particularly, the PSZs are much thinner, in many cases <1 cm (Sibson, 2003). Many fault zone studies, there-
fore, share a common denominator—only small amounts of sample are available. Such small amounts can
be delicate for the more classical geochemical and mineralogical analyses. In contrast, a relatively small
amount of material (< ~500 mg) with any shape (e.g., chip, powder, or drilling/cutting residue) is sufficient
for a complete set of magnetic analyses. Thus, in principle, magnetic changes at a small scale allow for resol-
ving the faulting‐related behavior. Magnetic scanning may reveal changes at the μm scale.
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Figure 4. Domain states of magnetite as a function of grain size and shapes, and examples for commonly used plots for
high‐field magnetic measurements. (a) Illustration of the domain state categories for magnetite at room temperature
and their relationship with shape and grain size of the magnetic particles (redrawn from Roberts et al., 2018).
(b) Example of a hysteresis loop with definition of M
rs
,M
s
,B
c
, and B
cr
, with the inset showing a backfield
demagnetization curve, with the definition of B
cr
. The high‐field slope (typically, B> 0.7 T) on the hysteresis loop is
defined as high‐field magnetic susceptibility (χ
hf
). The dark and gray lines are the hysteresis loops before and after
correction for the nonferrimagnetic matrix contribution, respectively. In this example the correction is a paramagnetic
correction as χ
hf
is positive. (c) Example of a Day plot of the hysteresis ratios M
rs
/M
s
and B
cr
/B
c
. Single domain (SD),
pseudo single domain (PSD), and multidomain (MD) boundaries are after Day et al. (1977). Note that recent Day plots
often use vortex rather than PSD following Roberts et al. (2018). The data points shown are host and fault zone
sediments from the frontal prism in Japan Trench cored by the Integrated Ocean Drilling Program Expedition 343, Japan
Trench Fast Drilling Project (JFAST) (reproduced from Yang et al., 2018). The host sediments lie mainly in the PSD
(or vortex) field, whereas most of the fault zone samples are located in the SD region. (d) Example of a FORC diagram
(first‐order reversal curve) for one of the fault zone sediments (FZ697) shown in Figure 4c. B
c
is equivalent to
particle coercivity, and B
u
to the local interaction field. Colors in the diagram represent absolute values of FORC density.
The B
c
peak centered at 30–40 mT with a prominent “kidney”shape toward higher coercivity suggests the
occurrence of SD pyrrhotite. FORC data are from Yang et al. (2018) and are reprocessed with the FORCinel package
(Harrison & Feinberg, 2008) with the VARIFORC (Egli, 2013) option used. VARIFORC smoothing parameters: vertical
ridge Sc0 = 4, horizontal smoothing factor Sc1 = 7, central ridge Sb0 = 3, vertical smoothing factor Sb1 = 7,
horizontal lambda λc= 0.1, and vertical lambda λb= 0.1. (e) Example of an isothermal remanent magnetization
(IRM) acquisition curve. After application of a stepwise increasing magnetic field, the remanent magnetization
increases until a maximum value is reached, which is termed saturation IRM (SIRM or M
rs
). (f) Example of unmixing
of an IRM acquisition curve to identify magnetic coercivity distributions using the MAX UnMix package (reprinted
from Maxbauer et al., 2016). The sample is an anoxic lake sediment from Baldeggersee, Switzerland, and three
coercivity components are identified: detrital magnetite‐low coercivity component (Component 1), biogenic magnetite
(Component 2), and oxidized magnetite (or hematite)—Component 3.
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Anisotropy of magnetic susceptibility (AMS) and/or paleomagnetic studies of fault rocks (see sections 3.5
and 3.6 for details) require oriented samples. Classically, such samples are cylinders of 25 mm dia-
meter × 22 mm height or 20 mm on‐a‐side cubes (known as standard‐size samples), representing a volume
deemed statistically representative. Due to the limited width of a fault zone, oriented sampling in a narrow
fault zone is a challenge, and “contamination”by the host rock adjacent to the target layer(s) may compli-
cate the interpretation. Also, bias may creep into AMS results due to the unconventionally small sample size
dictated by fault rock samples. Almqvist et al. (2020), for example, found that the mean magnetic susceptibil-
ity and magnetic anisotropy degree (see section 3.5 for definitions) are inversely proportional with sample
size for a set of pseudotachylyte cubes from western Jämtland (central Swedish Caledonides) ranging in
volume between ~0.2 and ~0.03 cm
3
.“Minicores”of 12 mm diameter × 11 mm long, cubic mini‐AMS sam-
ples with 1 cm edges, or even down to a size of 3.5 mm on‐a‐side (hence ~250 times volumetrically smaller
than standard‐size samples), often work fine. AMS analysis of pseudotachylyte fabric on such 3.5 mm cubes
enabled determining the focal mechanisms of ancient earthquakes (Ferré et al., 2015, 2016). Despite sample
size issues for very small samples where a cautious and careful interpretation is appropriate, AMS studies on
small samples are nonetheless a promising new avenue to unveil detailed geological features (Almqvist
et al., 2020; Ferré et al., 2015, 2016; Zhu et al., 2017).
Berndt et al. (2016) demonstrated that samples with a thermoremanent magnetic moment larger than
10
−11
Am
2
irrespective of their sizes (ranging from the centimeter scale to tens of nm scale) contain enough
magnetic particles to be accurate magnetic recorders. Paleomagnetic measurements can thus also be carried
out on smaller samples than the standard size samples. Millimeter‐sized samples (e.g., Böhnel et al., 2009;
Suttie et al., 2010) or even smaller samples of geological materials like single silicate crystals (e.g., zircon, pla-
gioclase, and olivine crystals) can yield accurate paleointensity and paleodirection estimates using ultrahigh
sensitivity moment magnetometry and advanced demagnetization techniques (e.g., Berndt et al., 2016; Fu
et al., 2017; Sato et al., 2015; Tarduno et al., 2015). These include, but are not limited to, scanning SQUID
(superconducting quantum interference device) microscopy (SSM; e.g., Weiss et al., 2007), quantum dia-
mond magnetometry (QDM; e.g., Glenn et al., 2017), and microwave demagnetization (e.g., Suttie
et al., 2010). This enables paleomagnetic analyses of the fault rocks. For such studies, oblique drilling across
a fault zone may offer a way to orient the samples by making use of the present‐day NRM overprint that is
present in most samples.
3.2. Field‐Dependent Measurements: Low‐Field Magnetic Susceptibility
One of the most widely used room temperature magnetic measurements is the low‐field magnetic suscept-
ibility (volume specific: κ
lf
, or mass specific: χ
lf
), measured in small applied magnetic fields, typically,
~200–300 A/m; that is, several times the strength of the Earth's magnetic field (~25–50 A/m). It is a measure
of the “magnetizability”of a sample and expresses in a general way the classes of magnetic materials (e.g.,
Dunlop & Özdemir, 1997): dominantly ferromagnetic or ferrimagnetic (very high χ
lf
), paramagnetic (small
positive χ
lf
), or diamagnetic (very small negative χ
lf
). A rock is an ensemble of ferromagnetic (sensu lato),
paramagnetic, and diamagnetic minerals. Therefore, to characterize the ferrimagnetic (χ
ferri
) component
(s) (i.e., magnetite, pyrrhotite, or greigite) in samples, contributions from paramagnetic (χ
para
), diamagnetic
(χ
dia
), and imperfect antiferromagnetic minerals (i.e., hematite and goethite) should be subtracted from the
bulk low‐field magnetic susceptibility (χ
lf
). The imperfect antiferromagnetic minerals are often included in
the paramagnetic contribution since their susceptibility does not differ that much. Magnetic susceptibility
can also be measured at different frequencies of the applied alternating current (AC) field (typically varying
between a few hundreds of Hz to a few kHz), yielding the frequency‐dependent magnetic susceptibility,
expressed either as a percentage of the low‐field susceptibility (κ
fd
%) or as a mass‐(χ
fd
) or volume‐specific
number (κ
fd
). With increasing measurement frequency, the SP/stable SD (SSD) boundary (~25 nm at room
temperature for a relaxation time of 100 s) shifts to smaller volumes. Hence, for a given grain‐size distribu-
tion more grains become blocked at a higher measurement frequency and are SD with a lower susceptibility
than SP particles, which explains the difference in susceptibility between the low and high measurement fre-
quency. Frequency‐dependent magnetic susceptibility is thus an effective way to determine the concentra-
tion of magnetic particles over a small grain size window across the SP/SSD boundary (e.g., Q. S. Liu
et al., 2012).
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The diamagnetic and paramagnetic magnetic moment increases linearly with applied field up to very high
field values; hence, their susceptibility is constant for a given temperature. Also, magnetite's magnetization
is linear with the applied magnetic field up to ~800 A/m, that is, the entire field range of low‐field suscept-
ometers. However, the magnetization of some other ferrimagnetic minerals, such as pyrrhotite, hematite,
and titanomagnetite, shows a marked applied field dependence already in the low field of interest here
(de Wall, 2000; Hrouda, 2011; Hrouda et al., 2006; Jackson et al., 1998; Worm et al., 1993). Variation of
the magnetic susceptibility as a function of applied field can thus diagnose the presence of these magnetic
minerals in fault zones. The bridge‐type susceptometers are the most versatile and widely used instruments
for susceptibility measurements. This includes field‐, frequency‐, and temperature‐dependent magnetic
susceptibility measurements, AMS measurements (see also section 3.5), and for some instruments, measure-
ment of the in‐phase and out‐of‐phase components of the magnetic susceptibility (e.g., Hrouda et al., 2016;
Hrouda & Ježek, 2014).
3.3. Field‐Dependent Measurements: High‐Field Magnetic Measurements
3.3.1. Hysteresis Loops and FORC Diagrams
Magnetic hysteresis loops (Figure 4b) are not only informative of the magnetic mineralogy but also indica-
tive of the dominant domain structure of magnetic minerals in the sample. Saturation magnetization (M
s
)is
the largest possible magnetization of a sample; it is thus a measure of the total amount of ferromagnetic
minerals in the sample. The saturation remanence (M
rs
) is the corresponding remanent magnetization after
removal of the applied field. The coercivity (B
c
) and remanent coercivity (B
cr
) are measures of magnetic sta-
bility. The ratios M
rs
/M
s
and B
cr
/B
c
are commonly plotted on the so‐called Day plot (Figure 4c; Day
et al., 1977; Dunlop, 2002a, 2002b; Dunlop & Özdemir, 1997; Lanci & Kent, 2003) as indicators of domain
states and, indirectly, grain size. As an alternative, the plot of M
rs
/M
s
versus B
c
, referred to as “squareness
plot”in the engineering literature, was proposed to avoid the need of determining B
cr
and to remove poten-
tial ambiguities associated with the B
cr
/B
c
ratio: Particles with markedly different B
c
values may have their
B
cr
/B
c
largely overlaying complicating interpretation of the domain state (Tauxe et al., 2002). The Day plot is
a classic rock magnetic tool and has been used extensively for domain state diagnosis since it was proposed
by Day et al. (1977). However, several fundamental ambiguities associated with its interpretation have been
documented recently by Roberts et al. (2018). Additional constraints, other than those offered by the Day
plot, are required to properly interpret the hysteresis parameters (i.e., dominant domain state) in terms of
physical grain size. Beyond the Day plot, the first‐order reversal curve (FORC) technique has been developed
to derive more information from hysteresis measurements (e.g., Roberts et al., 2000, 2006, 2014; Zhao
et al., 2015). FORC diagrams (an example in Figure 4d) provide important insights into not only the mag-
netic minerals but also the distribution of coercivity and magnetostatic interactions among magnetic parti-
cles (see review by Roberts et al., 2014, for more details). Given the interpretive ambiguities in the Day
diagram, unmixing of FORC diagrams (Lascu et al., 2015), and determination of remanent, transient, and
induced FORC diagrams (P. X. Hu et al., 2018; Zhao et al., 2017) have been suggested as suitable candidates
for diagnosing domain state on a component‐by‐component basis (Roberts et al., 2018). However, when
samples are dominantly paramagnetic or antiferromagnetic with only a small ferrimagnetic contribution,
the low signal‐to‐noise ratio prevents acquisition of meaningful FORC data, that is, the ferrimagnetic signal
is weak in comparison to the total magnetization (e.g., Jackson & Solheid, 2010). For those samples hyster-
esis loop measurement with determination of the classic hysteresis parameters is still the foremost way to
characterize their magnetic behavior. In addition, hysteresis measurements also allow for assessing the rela-
tive magnetic contribution of ferrimagnetic minerals and that of the matrix minerals (i.e., paramagnetic and/
or diamagnetic) through comparison of the high‐field magnetic susceptibility (χ
hf
, i.e., the slope of the linear
high‐field segments of the hysteresis loop, cf. Figure 4b) and the low‐field susceptibility (χ
lf
) of a sample
(Figure 4b).
3.3.2. Remanence Measurements at Room Temperature
High‐field methods are most commonly used to identify magnetic mineral phases. IRM acquisition curves
(Figure 4e) obtained at room temperature in successively increasing direct fields contain a wealth of infor-
mation about magnetic hardness. Unmixing of IRM acquisition curves along with forward basis functions
can provide information that is diagnostic of magnetic mineralogy (Figure 4f; Egli, 2004a, 2004b, 2004c;
Kruiver et al., 2001; Maxbauer et al., 2016). Collections of IRM acquisition curves can also be subjected to
end member modeling, an inverse bilinear unmixing approach (Heslop & Dillon, 2006; Weltje, 1997).
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The end members then serve as basis for a geological interpretation of the variability in the IRM acquisition
curves. It has been used for paleoenvironmental analysis, the detection of magnetotactic bacteria (e.g., Just
et al., 2012), and in the context of remagnetization, the (partial) resetting of the NRM in rocks at some point
during their geological history (e.g., Aben et al., 2014; Gong et al., 2009; Huang et al., 2015).
ARM is generated by placing a sample in a small steady direct current (DC) field, the DC bias field (of the
order of the Earth's magnetic field), and superimposing an AC field of which the amplitude is steadily
ramped down from a preset initial value to zero. It is a useful laboratory technique for characterizing
magnetic particles (Q. S. Liu et al., 2012). It senses SD magnetite particularly well. The different behavior
of ARM and χ
lf
with grain size make plots of ARM versus χ
lf
(i.e., the Banerjee diagram, Banerjee et al., 1981)
or χ
ARM
(χ
ARM
= ARM/DC bias field) versus χ
lf
(i.e., the King diagram, King et al., 1982) useful for detecting
grain size changes, in particular for magnetite.
Another rock magnetic parameter, which is often employed to measure the relative contributions of low and
high coercivity material to a sample's saturation isothermal remanent magnetization (SIRM), is the S‐ratio.
It compares the SIRM obtained in a saturating forward field with the IRM obtained subsequently in a suita-
ble backfield (usually a reversed field of 0.3 T, yielding IRM
−0.3T
). The classic S‐ratio, S
classic
,isdefined as
(e.g., Thompson & Oldfield, 1986) follows:
Sclassic ¼−IRM−0:3T=SIRM (1)
It scales from −1 (only high coercivity material) to 1 (only low coercivity material). It can also be defined
as (Bloemendal et al., 1992) follows:
S¼1−IRM−0:3T=SIRMðÞ=2 (2)
or
Sforward ¼IRM0:3T=SIRM (3)
where the IRM
0.3T
is remanent magnetization acquired in a forward field of 0.3 T. The S‐ratios in
Definitions (2) and (3)—the latter is also called “forward S‐ratio”(Heslop, 2009)—scale from 0 to 1; they
provide a measure of the relative contribution of the low coercivity material to the total remanence. It
should be realized that numeric values of Definition (1) on the one hand versus (2) and (3) on the other
are different and cannot be used interchangeably without the definitions being specified. When the S‐ratio
approaches unity, in general, ferrimagnetic minerals are dominant over antiferromagnetic minerals, and
vice versa. However, it is important to note that variations in the relative concentration of high‐and
low‐coercivity phases are nonlinear and interpretation of the S‐ratio is nonunique, because it is not solely
a function of mineral type but is also influenced by factors such as grain size and substitution. The S‐ratio
is an example of so‐called closed data (i.e., data within certain bounds like fractions or percentages); the
additive‐log‐ratio transform (e.g., Swan & Sandilands, 1995) has been suggested to obtain meaningful
descriptive statistics for the S‐ratio information (Heslop, 2009).
3.4. Temperature‐Dependent Magnetic Measurements
3.4.1. High‐Temperature Magnetic Measurements
High‐temperature magnetic analyses, either in‐field measurements or remanence measurements, are
among the most diagnostic measurements for assessing magnetic mineralogy because magnetization
decreases with increasing temperature and becomes 0 at and above the magnetic ordering temperature since
the magnetic mineral has become a paramagnet. The magnetic ordering temperature is named Curie tem-
perature (T
C
) for ferromagnet and ferrimagnet and Néel temperature (T
N
) for imperfect antiferromagnetic
materials (Dunlop & Özdemir, 1997). For example, magnetite and hematite have T
C
and T
N
values of
578°C (Figure 5a) and 675°C (Figure 5b), respectively (Dunlop & Özdemir, 1997). The magnetite‐
ulvöspinel (Fe
3
O
4
‐Fe
2
TiO
4
) solid solution series and the hematite‐ilmenite (α‐Fe
2
O
3
‐FeTiO
3
) solid solution
series show a marked range in ordering temperatures respectively from 578°C to −153°C (T
N
of ulvöspinel)
and from 675°C to −233°C (T
N
of ilmenite) (Dunlop & Özdemir, 1997). With increasing Ti substitution
ordering temperatures decrease approximately linearly. Goethite has a T
N
of ~120°C (Figure 5c), which
decreases below room temperature with increasing Al substitution (Q. S. Liu et al., 2006). Goethite can
accommodate up to ~32 mol% Al substitution with a corresponding magnetic ordering temperature of
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−58°C (Murad & Bowen, 1987). Maghemite has a T
C
of 645°C (Figure 5d; Özdemir & Banerjee, 1984), but it
often inverts to hematite before the T
C
is reached: Depending on grain size and amount of substitution
maghemite can invert to hematite at any temperature between 250°C and 1000°C (Dunlop &
Özdemir, 1997). A typical inversion range for fine‐grained maghemite is between 300°C and 400°C (e.g.,
Q. S. Liu et al., 2005). Monoclinic pyrrhotite has a T
C
of 325°C (Figure 5e); T
C
is slightly lower for
Figure 5. Examples of thermomagnetic curves of some common magnetic minerals, showing their high‐temperature
behavior including Curie or Néel temperatures. “Stepwise”thermomagnetic analyses were performed in air with a
modified horizontal translation Curie balance (Mullender et al., 1993) in Paleomagnetic Laboratory Fort Hoofddijk,
Utrecht University (The Netherlands). Details on samples, heating/cooling rates, temperature cycles, and applied fields
are given in each panel. Kadan hematite (panel b) is from the Czech Republic and Robe River maghemite (panel d)
is from Western Australia. The Hopkinson peak (panel c) is a manifestation of superparamagnetism: the coercivity of the
goethite is very high (tens of teslas at room temperature) in comparison of the field range used in the
thermomagnetic analysis. Then the magnetic moment can be considered a “low‐field”susceptibility which goes through
a maximum at the Néel/Curie temperature.
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Ni‐substituted pyrrhotite (Kontny et al., 2000; Schwarz & Vaughan, 1972). T
C
remains undetermined for
greigite because greigite decomposes before the T
C
can be measured (like the situation of maghemite): It
exceeds 400°C (Roberts et al., 2011). Diagnostic for greigite is its thermochemical alteration to a nonmag-
netic phase starting at ~200°C and finishing at ~350°C when heated in air (Figure 5f; Dekkers et al., 2000;
Reynolds et al., 1994; Roberts, 1995; Torii et al., 1996). At higher temperatures, above 400°C magnetite forms
which in turn is converted to hematite above 500–550°C yielding the typical “greigite‐in‐air”thermomag-
netic signature. To date, greigite has not been reported in fault rocks; however, discrimination between grei-
gite and pyrrhotite is essential for fault zone magnetic studies, as the latter is commonly present as product of
thermal alteration of pyrite in fault zones. In addition, with stepwise thermal demagnetization of three
orthogonal IRM components the thermal behavior of high‐, intermediate‐, and low‐coercivity magnetic
phases can be evaluated separately (Lowrie, 1990).
3.4.2. Low‐Temperature Magnetic Measurements
A major disadvantage of high‐temperature measurements is that the magnetic signal may be obscured by
newly formed magnetic mineral(s) during the thermal treatment. This applies to in‐field measurements in
particular. During measurements at low temperature, down to 4 K, thermochemical alteration is nonexistent
because the sample is not heated. Therefore, low‐temperature measurements are increasingly utilized in
rock magnetism (Dunlop & Özdemir, 1997; Q. S. Liu et al., 2012; Moskowitz et al., 1998; Rochette et al., 1990).
They are designed to determine magnetic ordering temperatures below room temperature and whether or
not so‐called low‐temperature transitions in magnetic minerals occur. For example, the well‐known
Verwey transition (T
V
) at ~120 K or −153°C (Figure 6a; Verwey, 1939) is diagnostic of magnetite, when it
undergoes a crystallographic phase transition from a cubic to a monoclinic structure. Similarly, the Morin
transition (T
M
,~250 K) at which hematite's magnetic structure undergoes a spin flop, a change in orienta-
tion with respect to its crystal structure, is a key indicator of the presence of (crystalline) hematite
(Figure 6b; Morin, 1950). Monoclinic pyrrhotite exhibits the Besnus transition at 30–34 K (Figure 6c;
Besnus & Meyer, 1964; Dekkers et al., 1989; Rochette et al., 1990, 2011). Goethite shows an increasing mag-
netization with decreasing temperature (Figure 6d; Guyodo et al., 2006; Rochette & Fillion, 1989). The
low‐temperature behavior with possible magnetic transitions is usually measured with a so‐called magnetic
properties measurement system (MPMS); both in‐field and remanent magnetizations are measured as a
function of applied field and temperature between 1.8 and 400 K. The most commonly used measurement
sequences and their advantages have been described in detail by Bilardello and Jackson (2013).
However, one must exercise caution when interpreting low‐temperature data to avoid ambiguities as some
minerals show similar features at the same temperature range. At low temperatures other Fe‐bearing mate-
rials, for example, Fe or Mn carbonates (siderite (FeCO
3
), and rhodochrosite (MnCO
3
)) order magnetically
at 37 and 34 K, respectively (Frederichs et al., 2003; Housen, Banerjee, & Moskowitz, 1996). Vivianite Fe
3
(PO
4
)
2
·8H
2
O) orders as well below 20 K (Frederichs et al., 2003). Fe‐bearing silicates (clay minerals, chlorite,
amphiboles, and pyroxenes) may also order below 15–20 K because of low thermal vibration at those low
temperatures. SP particles show distinctive low‐temperature behavior (e.g., Chang et al., 2009); spin glasses
may freeze in and iron oxide coatings around silicates may order below 50 K (e.g., Franke et al., 2007).
Therefore, low‐temperature measurements deliver very useful information, but their interpretation requires
expert knowledge. One should perform more elaborate measurement strategies (e.g., Bilardello &
Jackson, 2013; Kars et al., 2011) or combine the results with other magnetic and/or nonmagnetic measure-
ments. A nice example is the study of Kars et al. (2011). Through applying a magnetic field of 5 μT inside the
SQUID MPMS during the cooling of a room temperature SIRM to detect the presence of a magnetic ordering
temperature, they successfully diagnosed siderite and rhodochrosite in overmatured claystones from a bore-
hole in the Netherlands and in outcrop samples from the Borneo Prism; the previous interpretation of
fine‐grained pyrrhotite had to be abandoned.
3.5. Anisotropy of magnetic susceptibility and remanence
The magnetic anisotropy of rocks, that is, the magnetic properties of a given sample vary depending on the
sample orientation, results from the contributions of ferromagnetic (sensu lato), paramagnetic, and diamag-
netic minerals. Detailed coverage on the theoretical background, mineral sources, measurement procedures,
quantification, geological applications, and advantages and limitations of AMS and anisotropy of magnetic
remanence (AMR), hereafter collectively called “magnetic fabric,”can be found in several textbooks and
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review papers (Borradaile & Henry, 1997; Borradaile & Jackson, 2004, 2010; Ferré, Gébelin, et al., 2014;
Hirt, 2007; Hrouda, 1982; Jackson, 1991; Martín‐Hernández et al., 2004; Martín‐Hernández & Ferré, 2007;
Parés, 2015; Potter, 2004; Rochette et al., 1992; Tarling & Hrouda, 1993). Below we briefly outline the
rationale for AMS approaches.
AMS depicts geometrically the orientation and degree of alignment of the preferred orientations of mineral
grains, the mineral grain spatial distribution or their lattice‐preferred orientation, and the shape or
crystalline anisotropy of the grains in a rock (Figures 7a and 7b; e.g., Ferré, Gébelin, et al., 2014;
Martín‐Hernández & Ferré, 2007). It is commonly expressed by a symmetric second‐order tensor (Tarling
& Hrouda, 1993) and described by a triaxial ellipsoid with principal axis' lengths equal to the eigenvalues
of the magnetic susceptibility tensor, κ
max
≥κ
int
≥κ
min
, which correspond to the maximum, intermediate,
and minimum susceptibility axes, respectively (Figure 7b). The degree of anisotropy of the ellipsoid is mea-
sured by the parameter P(Nagata, 1961) or the “corrected”P,P
j
(Jelínek, 1981), which are respectively
defined as follows:
κmax=κmin (4)
and
Pj¼exp 2 lnκmax −lnκm
ðÞ
2þlnκint −lnκm
ðÞ
2þlnκmin −lnκm
ðÞ
2
1=2(5)
where κ
m
is the mean magnetic susceptibility, κ
m
=(κ
max
+κ
int
+κ
min
)/3. The shape of the ellipsoid is
characterized by the parameter T(Jelínek, 1981), which is defined as follows:
Figure 6. Low‐temperature treatment of (S)IRM for some common magnetic minerals, showing the low‐temperature
transition of magnetite (a), hematite (b), and monoclinic pyrrhotite (c). (a) Warming curve of SIRM acquired in 2.5 T at
5 K of a 37 nm sized stoichiometric magnetite showing the Verwey transition (T
V
) at ~120 K typical of magnetite
(Verwey, 1939). Data are from Özdemir et al. (1993). (b) Low‐temperature cycling of a 2.5 T SIRM at 300 K of a MD
hematite crystal (0.24 mm × 2.15 mm × 3.2 mm in size) from Mount Wright, Québec (Canada), showing its Morin
transition (T
M
) at ~250 K (Morin, 1950). Data are from Özdemir et al. (2008). (c) Low‐temperature cycling of SIRM
imparted at 300 K with a 4 T field for monoclinic pyrrhotite illustrating the Besnus transition (T
Bes
)at30–34 K
(Besnus & Meyer, 1964; Rochette & Fillion, 1988; Rochette et al., 1990). The sample with grain size of 250–150 μm was
separated magnetically from a pyrite/pyrrhotite skarn collected from Temperino in Tuscany, Italy (Dekkers, 1988;
low‐temperature data are from Dekkers et al., 1989). (d) Low‐temperature cycling (300–10–400–10 K) of an IRM acquired
at 300 K in a 2.5 T field for a synthetic goethite sample showing its low‐temperature magnetic behavior. Data are from
Guyodo et al. (2006).
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T¼2ln κint=κmin
ðÞ=ln κmax=κmin
ðÞ−1 (6)
Tranges over the spectrum of ellipsoid shapes (Figure 7c): 0 ≤T≤1 for oblate ellipsoids (disk shaped),
−1≤T≤0 for prolate ellipsoids (rod shaped), and T= 0 for neutral ellipsoids. Along with T, the magnetic
lineation L,
L¼κmax=κint (7)
and magnetic foliation F,
F¼κint=κmin (8)
are most informative to describe the AMS ellipsoid (Tarling & Hrouda, 1993). They are commonly plotted
on a so‐called Flinn diagram (Figure 7c; Flinn, 1965; Jelínek, 1981).
AMS can detect incipient deformation well before other strain indicators (e.g., Almqvist & Koyi, 2018;
Burmeister et al., 2009). The low‐field AMS, with assets of ease of data acquisition (minutes) and sensitivity
(down to one per mille), has been used almost exclusively in classic petrofabric analysis studies (e.g.,
Borradaile, 1988; Graham, 1966; Hrouda, 1982; Martín‐Hernández et al., 2004). In the context of fault
rocks—of special interest here—we note the extensive review by Ferré, Gébelin, et al. (2014) on magnetic
fabrics in ductile shear zones. In addition, measurement of AMS at high fields (e.g., Ferré et al., 2004;
Hrouda & Jelínek, 1990; Kelso et al., 2002; Martín‐Hernández & Ferré, 2007; Martín‐Hernández &
Hirt, 2001; Rochette et al., 1992), and temperature‐dependent AMS (Issachar et al., 2016, 2018;
Parés & van der Pluijm, 2014; Schmidt et al., 2007) have been proposed to separate the ferrimagnetic from
the nonferrimagnetic (i.e., paramagnetic and diamagnetic) contributions to the AMS. Recently,
frequency‐dependent AMS was proposed to determine both the presence of SP particles and the relative ten-
sor contributions of the SP and larger size fractions to the overall magnetic fabric (Hrouda & Ježek, 2014).
Figure 7. Principle of the AMS method for revealing petrofabrics. (a) A schematic diagram of a rock fabric. (b) The AMS
ellipsoid is composed of three orthogonal principal axes: κ
max
: maximum, κ
int
: intermediate, and κ
min
: minimum.
The κ
max
and κ
min
axes, respectively, correspond to long and short axes of grains or crystals. (c) Flinn diagram of the
lineation (L) versus foliation (F) illustrating the shapes of AMS ellipsoids in terms of the corrected degree of
anisotropy (P
j
, Tarling & Hrouda, 1993). The shape parameter (T) is also shown. Figures 7a and 7b are reprinted from
Cho et al. (2015), and Figure 7c is modified from Dubey (2014).
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This would help to evaluate the occurrence of grain fining of preexisting magnetic particles as a consequence
of the intensive shearing during earthquake slip (see section 4.2).
The measurement of magnetic remanence‐related anisotropies is much more time consuming and therefore
less utilized. It includes measurement of anisotropy of anhysteretic remanent magnetization (AARM) (e.g.,
Issachar et al., 2018; Jackson et al., 1991), anisotropy of isothermal remanent magnetization (AIRM) (e.g.,
Kodama & Dekkers, 2004; Stephenson et al., 1986), and anisotropy of thermal remanent magnetization
(ATRM) (Borradaile & Lagroix, 2000; Hirt, 2007). Such AMR data have been suggested to be an efficient
way to characterize the individual contributions of ferrimagnetic and antiferromagnetic minerals
(Martín‐Hernández & Ferré, 2007). Ferrimagnetic and paramagnetic/diamagnetic minerals in fault rocks
are usually formed at different faulting stages and/or through different faulting‐related process (e.g., fric-
tional heating and fluid‐rock interactions; see section 4). Those minerals, therefore, may reflect diverse
deformation paths. Examination of AMR across a fault zone might thus help diagnosing the subfabric
carried by the newly formed ferrimagnetic mineral(s) due to seismic frictional heating. Also, the thermal
remanent magnetization (TRM) imprint across fault zones may be evaluated, providing insights into the
dynamic deformation processes in a fault zone. The AMR techniques, including instrumental requirements,
measurement strategy, and pros and cons compared to standard low‐field AMS measurements, have been
discussed in detail by Martín‐Hernández and Ferré (2007).
3.6. Paleomagnetic Considerations
The NRM of rocks is constituted of one or several component(s) that as a rule of thumb represent(s)
different times in the geological history of that rock. The same pertains to fault rocks: The faulting gen-
erates frictional heat that (partially) resets the existing NRM. To properly isolate these NRM components,
stepwise alternating field or thermal demagnetization should be carried out with a sufficient number of
demagnetization steps according to standard paleomagnetic practice (e.g., Tauxe, 2010). The NRM demag-
netization behavior is visually inspected on vector endpoint plots (also referred to as Zijderveld plots;
Zijderveld, 1967), and the paleomagnetic directions are typically determined with principal component
analysis (Kirschvink, 1980). Different NRM components in fault rocks, especially the new so‐called
“secondary”component(s) formed during the rupturing itself (see section 4.7), may hold valuable
information on faulting‐related effects (e.g., Chou, Song, Aubourg, Lee, et al., 2012; Ferré, Geissman,
et al., 2014; Leibovitz, 2016).
3.7. Nonmagnetic Measurement Techniques
Fault rocks contain mixed magnetic mineral assemblages of potentially different origin, grain size, and
relative concentrations. Several types of magnetic measurements must therefore be used in concert to
properly identify magnetic carriers and minimize ambiguities. In some cases, the magnetic results are
supplemented by data acquired with what is referred to as “nonmagnetic techniques”to further unravel
their nature and causal links to faulting processes. The most frequently used nonmagnetic techniques
include, but are not limited to, the following: Mössbauer spectroscopy (e.g., Dyar et al., 2006), scanning
electron microscopy (e.g., Reed, 2005), transmission electron microscopy (e.g., McLaren, 2005), X‐ray
diffraction analysis (e.g., Lavina et al., 2014), electron probe microanalysis (e.g., Reed, 2005), magnetic
force microscopy (e.g., Grütter et al., 1995), electron spin resonance (e.g., Fukuchi, 2003; Pan &
Nilges, 2014), and diffuse reflectance spectroscopy (e.g., Torrent & Barrón, 2008). All these measure-
ments can provide important insights into the valency of Fe in minerals, preferred grain alignment,
grain size, and geochemistry of Fe‐bearing minerals, and help constraining the magneto‐mineralogical
interpretation.
4. Drivers for Changes in Magnetic Properties of Fault Rocks
As mentioned above, faulting is a complex process involving various physical and chemical changes, which
complicates the interpretation of fault rocks by any type of measurement, thus including the rock magnetic
approach reviewed here. The magnetic properties of fault rocks are determined by several mechanisms. In
this section we focus on those for which a consensus has been reached. These include (i) strain‐induced mag-
netic changes, (ii) grain‐size variations of preexisting magnetic minerals by grain fining, (iii) thermochemi-
cal reactions induced by frictional heating, (iv) coseismic frictional melting with pseudotachylyte formation,
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(v) chemical alteration and neomineralization due to fluid percolation, (vi) IRM acquisition due to EQL, and
(vii) remagnetization of the NRM due to physicochemical processes.
4.1. Strain‐Induced Magnetic Changes
Stress and related strain effects may have significant influence on magnetic properties of rocks. Stresses (dif-
ferential or even hydrostatic) can directly affect the behavior of ferrimagnetic mineral grains, especially
(titano‐)magnetite through the effect of magnetostriction (also known as piezomagnetism; e.g., Bezaeva
et al., 2015; Kapička, 1988, 1992; Kean et al., 1976). Strain may cause reorientation and/or internal deforma-
tion of ferrimagnetic grains (e.g., Borradaile, 1988; Till & Moskowitz, 2014). This would lead to changes in a
variety of magnetic properties, the most important of which are magnetic anisotropy, changes in remanence
orientation and intensity, and changes in bulk magnetic properties such as coercivity and magnetic suscept-
ibility (e.g., Bezaeva et al., 2015; Borradaile, 1996; Gattacceca et al., 2007; Gilder et al., 2004, 2006; Jackson
et al., 1993; Jiang et al., 2013; Kapička, 1988, 1992; Kapička et al., 2006; Kean et al., 1976; Louzada et al., 2010).
For example, up to 25% shortening due to axial compression in a set of synthetic magnetite‐bearing “calcite
sandstone”samples irreversibly increases their coercivity and magnetic anisotropy but decreases their mean
magnetic susceptibility and the remanence component parallel to the shortening axis (Jackson et al., 1993).
Brittle or semibrittle fault rocks generally form in the upper ~10–15 km of the Earth's crust. The vertical
effective stress at this depth interval is about 160–240 MPa, calculated from an average rock density
(2,600 kg m
−3
) under the assumption of hydrostatic pore fluid pressure. The stress sensitivity of magnetic
susceptibility for various rocks ranges roughly from 0.8 to 1.3 × 10
−3
MPa
−1
(unit: one over pressure per unit
area; e.g., Kapička, 1988; Kean et al., 1976). The stress may induce decreases in a rock's magnetic susceptibil-
ity ranging from 12% to 30%, if only the vertical effective stress is taken into consideration. In this context,
the potential stress effects on magnetic properties of fault rocks cannot be ignored, at least not in seismogenic
zones. It has been documented that changes in magnetization of stressed rocks due to local accumulation of
stress may induce transient anomalies in the local geomagnetic field prior to an earthquake, referred to as
piezomagnetism or tectonomagnetism (e.g., Sasai, 2001; Yamazaki, 2013).
However, an interpretation is complicated by many factors, such as variations in magnetic carriers (mag-
netic mineralogy, concentration, and grain size), ambient magnetic field, stress loading pattern (e.g., hydro-
static or uniaxial, shock or static, single or cycling), and duration (e.g., Gilder et al., 2006, 2018;
Kapička, 1992; Louzada et al., 2010; Volk & Feinberg, 2019). For example, pressure loading removes part
of the remanent magnetization of a rock when the pressure is applied in a zero ambient magnetic field, a
procedure conveniently termed pressure demagnetization (e.g., Bezaeva et al., 2010; Gattacceca et al., 2007;
Jiang et al., 2013; Louzada et al., 2010, 2011; Pozzi, 1975; Volk & Feinberg, 2019; Volk & Gilder, 2016). The
higher the pressure, the more remanence is removed. A detailed compilation of the available experimental
pressure demagnetization data on (titano‐)magnetite‐, (titano‐)hematite‐, and pyrrhotite‐bearing rocks is
provided in Louzada et al. (2011). In contrast, a rock acquires magnetic remanence when the (cycling)
pressure loading is taking place in a magnetic field; such remanence is known as pressure remanent mag-
netization or piezo‐remanent magnetization (Nagata, 1966). The higher the ambient magnetic field, the
more remanence is acquired for a given pressure (Nagata & Carleton, 1969). Also, with increasing Ti sub-
stitution, the magnetostriction coefficient of magnetite increases (Dunlop & Özdemir, 1997; Gilder & Le
Goff, 2008). Thus, the magnetic susceptibility and remanence of titanomagnetite are much more sensitive
to the stress regime than stoichiometric magnetite (Dunlop & Özdemir, 1997; Gilder & Le Goff, 2008).
Moreover, the effect of pressure on magnetic susceptibility generally decreases with decreasing magnetic
grain size; the acquisition of pressure remanent magnetization or demagnetization is quite efficient in
large MD grains in comparison to small SD particles (e.g., Gilder et al., 2006; Kean et al., 1976; Volk &
Feinberg, 2019).
4.2. Grain Fining of Preexisting Magnetic Particles
A fault core evolves by grain comminution consisting of early bulk fragmentation and late abrasion of
grains (e.g., Billi, 2005; Hattori & Yamamoto, 1999; Sammis & Ben‐Zion, 2008; Storti et al., 2007), reflecting
a progressive reduction in grain size. So, in fault rocks the intensive shearing also may cause widespread
splitting of ferromagnetic (sensu lato) grains (if present) into (much) finer material (Figure 3a). Magnetic
susceptibility of ferrimagnetic particles depends strongly on grain size when the grains are close to the
SP/SSD threshold with SP particles having a distinctly higher magnetic susceptibility than SSD particles
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(e.g., Walden et al., 1999). Thus, magnetic susceptibility may be elevated in
fault rocks, even when there is no change in the total concentration of
magnetic minerals. Measuring frequency‐dependent magnetic susceptibil-
ity and low‐temperature magnetic measurements are powerful means for
estimating the SP particle contribution (e.g., Banerjee et al., 1993; Q. S. Liu
et al., 2012). Effects of grain fining as a consequence of shearing can thus
be evaluated by these two approaches. The identification of comminuted
nanograined magnetic particles is also informative on fracture energetics
of fault gouge, as calculation of total grain surface area is a common
approach to estimate the fracture energy associated with fault gouge for-
mation (e.g., Aretusini et al., 2017; Chou, Song, Tsao, et al., 2014; Wilson
et al., 2005).
4.3. Thermochemical Reactions Induced by Frictional Heating
The transient high temperatures involved in ruptures induce thermoche-
mical reactions within the fault zone (Figure 3a), including decomposi-
tion and transformation of Fe‐bearing minerals, particularly in the PSZ
(e.g., Han et al., 2007; Pei, Zhou, et al., 2014; Tanikawa et al., 2007,
2008; Yang et al., 2018, 2019). Fe‐bearing minerals that are widely present
in fault zones are siderite (FeCO
3
), lepidocrocite (γ‐FeOOH), goethite
(α‐FeOOH), pyrite (FeS
2
), smectite (Ca, Na, H) (Al, Mg, Fe, Zn)
2
(Si,
Al)
4
O
10
(OH)
2
‐xH
2
O) and chlorite ((Mg, Fe, Li)
6
AlSi
3
O
10
(OH)
8
). Most
relevant thermal reactions are described below.
1. Siderite's thermal decomposition product is magnetite (Koziol, 2004;
Pan et al., 2000) according to:
3FeCO3þ0:5O2
→
400–580°C Fe3O4þ3CO2(9)
To illustrate this reaction, high‐velocity friction experiments (to simulate seismic slip in the laboratory) on
siderite‐bearing fault gouge by Han et al. (2007) are insightful: they revealed that siderite was thermally
decomposed to nanocrystalline magnetite. This changed the gouge color from pale brown to black and
increased the magnetic susceptibility by nearly 3 orders of magnitude (Figure 8). It thus leaves a prominent
magnetic imprint of seismic slip in a fault zone, even when siderite is present only in minor amounts (Han
et al., 2007; Tanikawa et al., 2008).
2. Lepidocrocite is one of the four iron oxyhydroxides known in nature. It is a paramagnet at ambient
temperature and commonly occurs in fault gouges developed from granitic rocks (Fukuchi et al., 2005).
On annealing it converts to antiferromagnetic hematite (α‐Fe
2
O
3
) with strongly ferrimagnetic maghe-
mite (γ‐Fe
2
O
3
) as an intermediate phase (Fukuchi, 2003; Gendler et al., 2005). The conversion of lepi-
docrocite to maghemite (γ‐Fe
2
O
3
) starts at ~200°C, and the thermal dehydroxilation is completed at
~300–350°C, followed by inversion of maghemite to hematite (Gehring & Hofmeister, 1994; Gendler
et al., 2005). Maghemite is stabilized by its surface energy, so it is often prominently present as fine
particles (T. Chen et al., 2005). Particularly maghemite leaves a magnetic imprint, due to its ferrimag-
netic nature.
2γ−FeOOH
→
~200–350°C
γ−Fe2O3þH2O→γ−Fe2O3
→
> 500°C
α−Fe2O3(10)
3. Goethite, another iron oxyhydroxide, thermodynamically the most stable under ambient conditions,
also commonly occurs in fault zones. It dehydroxilates to hematite with traces of magnetite or
magnetite/hematite aggregates between 240°C and 400°C (Dekkers, 1990; Gualtieri & Venturelli, 1999;
Figure 8. Results of a high‐velocity friction experiment on a simulated fault
gouge consisting of a 1:1:1 mixture of siderite, quartz, and calcite
under a normal stress of 1.3 MPa and at a slip rate of 2.0 m/s conducted by
Han et al. (2007). (a) The sheared gouge shows markedly lower X‐ray
diffraction (XRD) peak intensities in comparison to the original
siderite‐bearing gouge. The X‐ray diffractogram also shows new magnetite
peaks on the sheared gouge. The gouge color changes from (b) pale
brown before the experiment to (c) black after the experiment, with an
increase in magnetic susceptibility (χ) of a few orders magnitude,
supporting the XRD interpretation of magnetite.
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Özdemir & Dunlop, 2000; Yang et al., 2019). The magnetite in particular can be sensed readily by rock
magnetic techniques.
2α−FeOOH
→
~240–400°C
α−Fe2O3þγ−Fe2O3=Fe3O4
ðÞ
þH2O (11)
4. Pyrite is often found as accessory mineral in faults developed within siltstones (e.g., Chou, Song,
Aubourg, Song, et al., 2012) and other sediments (e.g., Yang et al., 2018). It decomposes to pyrrhotite
(FeS
x
,1≤x< 1.08) as a result of heating to temperatures >500°C under reducing conditions that prevail
in the subsurface (Bhargava et al., 2009; Toulmin & Barton, 1964). Pyrrhotite is strongly ferrimagnetic
and its occurrence makes a magnetic imprint.
FeS2↔FeSxþ1−0:5xðÞS2(12)
5. Smectite and chlorite are among the most common constituents of fault gouge. On heating they desorb
iron which is precipitated as magnetite. Fine‐grained magnetite is produced when heating smectite over
250°C in the laboratory; nucleation and growth of the magnetite grains proceed up to ~450–500°C (Hirt
et al., 1993; Yang et al., 2019). Thermochemical reactions of the iron adsorbed to chlorite (particularly
chamosite) result in the neoformation of magnetite on heating to 400–700°C (Tanikawa et al., 2008).
Newly formed ferrimagnetic minerals (i.e., magnetite, maghemite, and pyrrhotite) as consequence of the
aforementioned thermochemical reactions may increase magnetic susceptibility and magnetization of fault
rocks up to a few orders of magnitude (e.g., Han et al., 2007; Mishima et al., 2006, 2009; Tanikawa et al., 2007,
2008; Yang et al., 2012a, 2018, 2019). These newly formed magnetic minerals can be readily diagnosed by the
different rock magnetic measurements (or combinations thereof) discussed in section 3. It is necessary to
exclude the possibility that the elevated magnetic susceptibility is caused by fining of ferrimagnetic particles
before attributing it to the magnetomineralogical changes. However, frictional heating is essentially con-
strained to the PSZ itself: The temperature rise induced by frictional heating decays exponentially away from
the PSZ, as widely reported in modeling results (e.g., Fulton et al., 2013; Yang et al., 2018; Yang, Chen,
et al., 2013). Therefore, frictional heating acts prominently only on a very narrow (millimeters to centi-
meters) zone enclosing the PSZ. Its effects attenuate quickly in other parts of a fault zone, for example,
the damage zone.
4.4. Coseismic Frictional Melting With Pseudotachylyte Formation
As mentioned earlier, pseudotachylytes may occur within seismic slip zones (e.g., Ferré et al., 2012, 2017;
Hirono, Ikehara, et al., 2006; A. Lin, 2008; A. Lin et al., 2005; Otsuki et al., 2009; Rowe & Griffith, 2015;
Sibson & Toy, 2006), due to frictional heating that raises the temperature for a short time above the rock's
melting point. The oxygen fugacity in the highly localized melt pockets is generally 2–4 log units above
the fayalite‐magnetite‐quartz buffer (~−11 log units at ~1000°C and 110–160 MPa, e.g., O'Hara &
Huggins, 2005). Pseudotachylytes are often more magnetic than their parent rocks, due to the presence of
fine‐grained PSD/SD ferrimagnetic minerals, predominantly magnetite, that form due to oxidation of mafic
silicates during the melting and subsequent quenching (Ferré et al., 2005, 2012, 2017; Ferré, Geissman,
et al., 2014; Nakamura et al., 2002; O'Hara, 2001; Pittarello et al., 2012; Zhang et al., 2018). This enables a
potentially excellent recording of the NRM in the pseudotachylytes during cooling through the blocking
temperature, that is, the NRM is a TRM and should be considered a primary remanence, representing the
time of pseudotachylytes cooling, that is, shortly after the rupture itself. In such a context, the paleomagnetic
baked contact test and examination of the additivity of partial TRM (e.g., Tauxe, 2010), which reveals the
unblocking temperature spectrum and its gradient across a fault zone, would be a means to locate the zones
with the highest temperatures in the pseudotachylyte vein zones. Recently, Zhang et al. (2018) found metal-
lic iron in vacuum fusion experiments (in argon) on cataclasites from Borehole No. 2 of the WFSD project. It
was formed by the reducing action of Fe‐bearing minerals above 1300°C. They proposed that the presence of
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metallic iron could be another reason for the magnetic highs of pseudotachylytes. Similarly, studies on
high‐pressure high‐temperature effects in other disciplines, such as impact glasses (e.g., Rochette et al.,
2015, 2019), fulgurite (e.g., Essene & Fisher, 1986), shocked Martian meteorites (e.g., Van de Moortèle
et al., 2007), have also documented that metallic iron was formed from reduction of more oxidized material.
It is hard to identify the metal nanoparticles with conventional mineralogical techniques (e.g., scanning
electron microscopy [SEM], Raman spectrometry, and Mössbauer spectroscopy; cf. Van de Moortèle
et al., 2007), the presence of metallic iron in natural pseudotachylytes may thus have been overlooked in
previous studies.
4.5. Chemical Alteration Due to Fluid Percolation
Fault rocks can be susceptible to disaggregation and dissolution by fluids during the coseismic, postseismic,
and interseismic periods (cf. section 2 and Figure 3). The short‐duration pulses of hot fluid induced by fric-
tional heating during the seismic slip (Ishikawa et al., 2008; Muir‐Wood, 1994) can easily dissolve preexisting
Fe‐bearing minerals (e.g., pyrite, siderite, and Fe‐silicates) whereby Fe
2+
is liberated (Humbert et al., 2012;
Kopp & Humayun, 2003) (Figure 3a). This is followed by precipitation of (among others) ferrimagnetic
minerals during the postseismic period because of the low solubility of iron (oxy)(hydr)oxides (Figure 3b).
During interseismic periods that last for very long times, fluids, both of meteoric and connate origin infiltrate
the fault zone and percolate along a dense microcrack and macrocrack network (Figure 3c). Fluid circula-
tion would lead to destabilization of iron‐bearing (clay) minerals in fault zones, with release of Fe
2+
that dis-
solves in the fluid and is subsequently transported (Grosz et al., 2006; Hashimoto & Kaji, 2012; Just &
Kontny, 2012). The Fe
2+
‐rich fluids thus are potential sources to precipitate magnetic minerals (Pechersky
& Genshaft, 2001). Upon reaction with an Fe
2+
‐bearing fluid at temperatures up to 250°C and the most com-
mon pH ranges (i.e., pH = ~4.5 at 250°C to ~6.5 at 50°C; Ohmoto, 2003), for example, hematite will trans-
form to magnetite. Magnetite in turn can be transformed to hematite in acidic solutions, with ferrous iron
getting into solution (i.e., a so‐called nonredox reaction, cf. Ohmoto, 2003):
α−Fe2O3þH2O fluid
ðÞ
þFe2þ↔Fe3O4þ2Hþ(13)
Reaction of lepidocrocite with dissolved Fe
2+
can form magnetite in alkaline media (Cornell &
Schwertmann, 2003):
2γ−FeOOH þFe2þ→Fe3O4þ2Hþ(14)
Similarly, the breakdown of pyrite produces goethite or hematite depending on temperature and pH. High
temperature and low pH generally favor hematite over goethite (Cornell & Schwertmann, 2003): at low tem-
perature and high pH (slightly alkaline conditions)
4FeS2þ15O2þ10H2O fluidðÞ→4α−FeOOH þ8H2SO4(15)
or at high temperature and low pH (moderately acidic conditions)
4FeS2þ15O2þ8H2O fluidðÞ→2α−Fe2O3þ8H2SO4(16)
Magnetic properties of fault zones thus change drastically on (partial) removal of preexisting magnetic
minerals and/or formation of new (magnetic) minerals (e.g., Chou, Song, Aubourg, Song, et al., 2012;
Yang, Chen, et al., 2013; Yang, Yang, et al., 2016). However, the migration paths and composition of fault
fluids may vary enormously over time and space, as fault zones act as fluid‐flow conduits or barriers at dif-
ferent stages in their development (Goddard & Evans, 1995). This leads to temporal and spatial changes in
redox state of a fault zone as function of episodic fault rupture and healing cycles during the lifetime of
the fault (Yamaguchi et al., 2011). Thus, precipitation products of fluid circulation (e.g., iron oxides, iron
hydroxides, and/or sulfides) are highly dependent not only on the physical and chemical conditions (perme-
ability, temperature, pressure, redox and pH, etc.) and chemical composition of the fluids but also on the
fluid/rock ratio and the lithology of the fault zones (Chou, Song, Aubourg, Song, et al., 2012; Grosz
et al., 2006; Guichet et al., 2006; Yang, Chen, et al., 2013; Yang, Yang, et al., 2016).
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4.6. EQL‐Induced IRM Acquisition
EQL currents have been reported on several occasions during large earthquakes (e.g., Fidani, 2010;
Heraud & Lira, 2011; Kamogawa et al., 2005; J. Y. Liu et al., 2015; St‐Laurent et al., 2006; Thériault
et al., 2014). However, the underlying formation mechanism is still controversial. A flurry of mechanisms
have been put forward to explain the charge generation in the subsurface during earthquake motion: piezo-
magnetism, piezoelectricity, electrokinetic flow processes, opening of the tips of cracks, stress‐activation of
positive holes, and hydromechanical rupture (Enomoto et al., 2017; Freund et al., 2007; Losseva &
Nemchinov, 2005; St‐Laurent et al., 2006). Regardless of its formation mechanism, the coseismic electric
current is expected to travel along the fault plane, which often acts as a good electrical conductor with a
conductivity of up to 4 orders higher than the surrounding insulating host rocks (Ferré et al., 2005).
According to the Biot‐Savart law, a transient strong azimuthal magnetic field will be produced perpendicu-
lar to the fault plane (Ferré et al., 2005, 2012; Losseva & Nemchinov, 2005). In this arrangement—lamellar
conductive rocks enveloped by insulating margins—fault gouge, pseudotachylite generation veins, parallel
or near parallel to the fault plane, will readily acquire an IRM. This IRM is a secondary NRM termed
“earthquake lightning‐induced remanent magnetization (EQLIRM)”(Enomoto et al., 2001; Enomoto &
Zheng, 1998; Ferré et al., 2005, 2012). EQLIRM is distinct from lightning NRM in rocks struck by surface
lighting; in the latter the lightning IRM remagnetization pattern is circular (either clockwise or counter-
clockwise) centered on the lightning‐bolt impact location, coupled with a decrease in magnetization inten-
sity away from the electric bolt (e.g., Losseva & Nemchinov, 2005; Verrier & Rochette, 2002). In the context
of EQLIRM fault rocks, several paleointensity tests (e.g., Ferré, Geissman, et al., 2014; Gattacceca &
Rochette, 2004) and normalization of the NRM intensity by M
s
are helpful to discriminate high NRM
caused by a strong ambient magnetic field from that caused by high concentrations of magnetic minerals.
If the latter can be excluded, the high NRM then may be attributed to a strong magnetic field induced by
seismically generated electrical currents.
4.7. Secondary NRM Acquisition Due to Physicochemical Processes in Fault Zones
The EQLIRM dealt with in the previous section is a rather special case. In general, neoformed magnetic
minerals (resulting from any thermochemical reaction and/or fluid‐rock interaction) may carry a chemical
remanent magnetization (CRM). Also, the mere cooling of a slip zone (hot coseismic fluids included) would
lead to an imprint of a (partial) TRM in a fault zone (Figure 9). These newly imprinted NRM components in
and near the slip zone are oriented parallel to the contemporaneous local geomagnetic field (Chou, Song,
Aubourg, Lee, et al., 2012). Evidently, a NRM direction will be different in a fault zone when the (latest or
most intense) rupture occurred any time later than the formation of the original rocks. The fault gouge
can retain the magnetic record during interseismic periods. However, these newly imprinted NRM compo-
nents may be (partially) overprinted again during later seismic events (Figure 9; see also Chou, Song,
Aubourg, Lee, et al., 2012).
Often, several of the aforementioned mechanisms act simultaneously; it is necessary to find support for indi-
vidual mechanisms and/or their acting in concert to better understand the magnetic changes within a fault
zone. For instance, the removal or decrease in concentration of certain magnetic mineral(s), and grain fining
of magnetic particles that might be induced by the faulting, can readily be evaluated by concentration‐
dependent magnetic properties and grain‐size‐dependent properties, respectively. This is a foremost asset
of rock magnetic techniques.
5. Magnetic Properties and Faulting Processes
5.1. Indicator of Seismic Slip Records
Proper slip zone identification is crucial to understand earthquake generation processes and paleoseismol-
ogy (e.g., Rowe & Griffith, 2015). Estimating earthquake recurrence rates requires proper identification of
slip zones that accommodated historic earthquake events. Also, the most intense slip conditions may be
revealed. These are fundamental ingredients for seismic risk assessment. However, to find robust evidence
of seismic slip is not straightforward from a practical point of view. It is to be expected that the aforemen-
tioned faulting‐related magnetic changes should leave a trace in the rock record, thus making magnetic
property analysis a potential indicator of seismic slip.
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5.1.1. Magnetic Locating of Pseudotachylyte Veins
Classically, pseudotachylytes have been considered as the only definitive evidence of ancient seismic activity
(Cowan, 1999; Rowe & Griffith, 2015). They are well known as “earthquake fossils”and less prone to natural
alteration on geological timescales (A. Lin, 2008). A strong NRM is often observed in pseudotachylyte veins
due to the presence of abundant PSD/SD magnetite (Figure 10; Ferré et al., 2005). This makes (paleo)mag-
netic measurements a robust way to locate pseudotachylyte veins, providing further linkage to past seismic
events, such as the kinematic solution (see section 5.4.1 for details). In addition, it is believed that magnetic
properties of pseudotachylytes are controlled by oxygen fugacity and thus vary systematically with depth of
formation (Ferré et al., 2012, 2017; Ferré, Geissman, et al., 2014; O'Hara & Huggins, 2005). Pseudotachylytes
that formed at shallower depths (<10–12 km) commonly exhibit coseismically formed maghemite or hema-
tite (e.g., Fukuchi, 2003; Petrik et al., 2003), while those formed at intermediate crustal depths (~20 km) are
generally dominated by magnetite (e.g., Davidson et al., 2003; Ferré, Geissman, et al., 2014; Nakamura &
Nagahama, 2001). In contrast, pseudotachylytes formed at even greater depths tend to be ilmenite‐domi-
nated (e.g., Ferré, Geissman, et al., 2014; Moecher & Steltenpohl, 2009). Therefore, rock magnetic properties
are informative on the ambient conditions of the earthquake rupture (Ferré et al., 2017; Ferré, Geissman,
et al., 2014).
5.1.2. Coseismic/Postseismic Neoformed Magnetic Minerals
Coseismic and/or postseismic neoformed magnetic minerals along with their related anomalies in magnetic
properties have also been widely reported during the recent “rapid response”fault zone drilling campaigns
after major earthquakes. These include the Taiwan Chelungpu Fault Drilling Project (TCDP) following the
1999 Chi‐Chi earthquake (e.g., Hirono, Lin, et al., 2006; Kano et al., 2006), the WFSD drillings following the
2008 Wenchuan earthquake (e.g., H. Li et al., 2015; Pei, Li, et al., 2014; Wang et al., 2014; Zhang et al., 2017),
and the Japan Trench Fast Drilling Project (JFAST) following the 2011 Tohoku‐oki earthquake
(e.g., Brodsky et al., 2020; Fulton et al., 2013; W. Lin et al., 2013; Yang et al., 2018). In TCDP Hole‐B, the
Fault Zone FZB1136, the PSZ of the 1999 Chi‐Chi earthquake was estimated to contain up to ~200 ppmv
magnetite from the magnetic susceptibility and saturation IRM peaks (Figure 11; Chou, Song, Aubourg,
Song, et al., 2012; Chou, Song, Aubourg, et al., 2014). One should bear in mind that the parameters used here
(i.e., magnetic susceptibility and saturation IRM) are dependent on both magnetic particle size and concen-
tration of magnetic minerals (e.g., Dunlop & Özdemir, 1997; Peters & Dekkers, 2003; Walden et al., 1999;
Figure 9. A conceptual model showing the magnetic recording cycle in fault gouge of the Fault Zone FZB1136, TCDP Hole‐B, which accommodated the Taiwan
1999 Chi‐Chi earthquake (reprinted from Chou, Song, Aubourg, Song, et al., 2012). (a) During an interseismic period, in the fault gouge the magnetic record
of the last earthquake is preserved through geological time. (b) During a coseismic period, the elevated temperature and chemical degradation lead toa
partial‐to‐complete demagnetization of the preexisting magnetic record in the slip zone and the “baked contact”zone. (c) During the subsequent postseismic
period, cooling of the slip zone and/or hot coseismic fluids leads to an imprint of TRM, while the neoformed magnetic minerals resulting from any form of
chemical process may carry a CRM. The newly imprinted (T)CRM component(s) thus leave(s) a magnetic record of the recent earthquake event.
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see also section 3). Therefore, one has to take into account that estimates of magnetic mineral concentrations
based on these two parameters yield a range of possible concentrations rather than a single value (evidently
with an uncertainty envelope). In contrast, usage of saturation magnetization (M
s
) to this end leads to a
single number for the magnetic mineral concentrations in a given fault rock sample because M
s
is only
proportional to the concentration of magnetic minerals in a sample. In Hole WFSD‐1, magnetic
susceptibility peaks occur at the fault zone FZ590 which was inferred to have accommodated the 2008
Wenchuan earthquake. The peaks were credited to newly formed magnetite due to the thermal
decomposition of Fe‐bearing clays (e.g., smectite and chlorite) by frictional heating during seismic slip
(Pei, Li, et al., 2014; Yang et al., 2012a). In the JFAST drill hole, partial thermal alteration of pyrite to
pyrrhotite was exclusively identified in three slip zones developed in the frontal prism sediments (Yang
et al., 2018). The reactions were induced by seismic frictional heating and related coseismic hot fluids
(Figure 12), thus intimately associated with earthquake events (Yang et al., 2018). Examination of the
magnetic mineral assemblage in fault zones may therefore be a prospective way for identifying the most
recent PSZ (Cai et al., 2019; Chou, Song, Aubourg, et al., 2014; Chou, Song, Aubourg, Song, et al., 2012;
Yang et al., 2012a, 2018, 2019).
5.1.3. Seismically Imprinted NRM Components
As mentioned in section 4.7, newly imprinted NRM component(s), that is, TRM and/or CRM, may also leave
a (paleo)magnetic record of recent seismic slip. For instance, in the aforementioned slip zone TCDP
Figure 10. (a) Image of a slab sample across a generation pseudotachylyte vein from the Santa Rosa Mountains,
California and (b) the spatial variation in the magnetic susceptibility of the slab. The map was produced by interpolation
over a grid of 170 susceptibility measurements performed with a Bartington MS‐2F probe with a sensitivity of
2×10
−6
SI. Each 10 mm a data point was collected. (c) NRM intensity along the X‐X′profile in Figure 10b. (d) Hysteresis
properties of pseudotachylyte veins plotted on a Day plot. Most specimens display hysteresis behavior consistent
with PSD grains. (e) NRM thermal decay curves of the two samples (Nos. 03 and 07) obtained from the pseudotachylyte
veins (Figure 10a), indicating low‐Ti magnetite as the principal remanence carrier. Figures 10a–10d and Figure 10e
are reproduced from Ferré et al. (2005) and Ferré, Geissman, et al. (2014), respectively.
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FZB1136, a stable single NRM component with a direction very close to the present‐day Earth's field was
identified within an ~16 cm thick fault gouge zone (Figure 13; Chou, Song, Aubourg, Lee, et al., 2012). It
is therefore inferred to be the coseismic paleomagnetic record of the recent 1999 Chi‐Chi earthquake
(Chou, Song, Aubourg, Lee, et al., 2012). In contrast, NRM records of two other slip zones (i.e., FZB1194
Figure 11. Magnetic susceptibility (a), S‐ratio (b), and estimated concentration profiles of magnetite and goethite
(c) in the Fault Zone FZB1136, TCDP Hole‐B (Taiwan). The PSZ is characterized by a peak in the magnetic susceptibility
(a) and the S‐ratio (b), which is consistent with a large amount of magnetite (c). The concentrations of magnetite and
goethite are estimated from the magnetic susceptibility and SIRM of the core samples using a quantitative model
(for details on the modeling strategy, see Chou, Song, Aubourg, et al., 2014). The lowest value of the S‐ratio is located
near the center of the fault gouge, corresponding to the highest concentration in goethite (c). (d) Magnetic hysteresis
loops and (e) NRM thermal decay curves for selected samples from the Fault Zone FZB1136. Within the gouge (e.g., the
sample at 1136.34 m), the principal maximum unblocking temperature (T
b
) of ~120°C is consistent with the Néel tem-
perature of goethite, while in the PSZ (e.g., sample at 1136.38 m), the maximum T
b
is close to 580°C, indicating the
presence of magnetite as main magnetic carrier. The latter concurs with the open hysteresis loops saturated below 0.3 T
in Figure 11d. These observations suggest that the magnetic susceptibility anomaly at the Fault Zone FZB1136
(Figure 11a) is attributed to the formation of magnetite, as a result of the thermal decomposition (>400°C) of para-
magnetic minerals (e.g., pyrite, siderite, and chlorite) during the 1999 Chi‐Chi earthquake. Figures 11a and 11d are
adapted from Hirono et al. (2007) and Mishima et al. (2009), respectively; Figure 11e is compiled from Chou, Song,
Aubourg, Song, et al. (2012); data in Figures 11b and 11c are from Chou, Song, Aubourg, Song, et al. (2012) and Chou,
Song, Aubourg, et al. (2014), respectively.
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and FZB1243) show directions far away from the present‐day field that were probably acquired before the
Brunhes‐Matuyama reversal (0.78 Ma). They thus have recorded ancient earthquake events and were not
reactivated during the 1999 Chi‐Chi earthquake; otherwise, these slip zones would have been overprinted,
at least partially (Chou, Song, Aubourg, Lee, et al., 2012).
5.1.4. Earthquake‐Induced AMS Fabric
AMS analysis of clastic dikes (see Appendix A for definition) has also successfully identified the unique mag-
netic fabrics of dikes emplaced by injection due to seismically triggered fluidization. In the Dead Sea Basin
(Israel), one of the most active tectonic areas in the Middle East, for instance, host sediments appear to be
characterized by oblate AMS ellipsoids with vertical axes of the minimum magnetic susceptibility (κ
min
) that
indicate a normal sedimentary magnetic fabric, whereas the liquefied sediments are characterized by triaxial
fabrics with subvertical κ
int
axes and subhorizontal κ
max
axes parallel to the dike strike (Levi et al., 2006a).
Similar AMS fabrics have been reported by Jayangondaperumal et al. (2010) in the Himalayan frontal thrust
(in the western Himalaya), by Lakshmi et al. (2017) in the Dauki fault (Shillong Plateau, India), and by Cho
et al. (2017) in the Dadaepo Basin (SE Korea). These contrasting magnetic fabrics have provided crucial evi-
dence for seismically triggered fluidization of clastic materials. Furthermore, Levi et al. (2018) recently pro-
posed a plot of the magnetic lineation versus the shape parameter of the AMS ellipsoid to distinguish
different seismites formed in soft sediments, that is, damage zones, gouges, earthquake‐triggered folds,
Figure 12. A cartoon showing the thermal alteration of preexisting pyrite with neoformation of pyrrhotite as a result of
seismic frictional heating and subsequent coseismic hydrothermal fluids in a pyrite‐bearing fault zone during an
earthquake (reprinted from Yang et al., 2018). During the seismic slip, (a) pyrite in the slip zone thermally decomposes to
pyrrhotite once the temperature is over ~640°C and (b) a pyrrhotite rim surrounding the pyrite core is formed
indicating an incomplete reaction. (c) Leaching by coseismic hot fluids, pyrite is altered and an intergrown rim of
pyrrhotite and barite is formed. (d) Precipitation of pyrite/pyrrhotite around the silicates and/or nucleation of pyrrhotite
on feldspar grains occurs at expense of iron and sulfur released by destabilization of pyrite and other minerals in fault
zone, during cooling of the coseismic hot fluids. The reader is referred to Yang et al. (2018) for more details.
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and breccia layers. It enabled categorizing seismites that underwent different deformation processes into
distinct groups in the above‐mentioned Dead Sea fault system (cf. Figure 4 in Levi et al., 2018). This
illustrates that AMS can be also used as a potential petrofabric tool for recovering paleoseismic records in
soft sediments (Levi et al., 2006a, 2006b, 2014, 2018).
5.2. Thermal History of Seismic Slip
Estimation of the temperature rise induced by frictional heating during an earthquake provides crucial
information on faulting weakening mechanisms (e.g., frictional melting, thermal decomposition of carbo-
nates, and thermal pressurization), the prevailing dynamic shear stress, slip velocity, and displacement
(e.g., Di Toro et al., 2009; Scholz, 2019; Sibson, 1977; Yao et al., 2016). At the same time, earthquake
energy budgets as a whole can be assessed (e.g., Scholz, 2019). Here, temperature information derived
from melt thermodynamics from experimental igneous petrology data, and pyrometamorphism in prin-
ciple could be very useful. The approach is widely used to quantitatively determine melting and crystal-
lization temperatures (e.g., Berman et al., 1995; Ghiorso & Gualda, 2015; Palmer et al., 2015). However,
the minerals (e.g., olivine, clinopyroxene, biotite, and garnet) typically used as “temperature proxy”are
often not occurring in fault rocks, in particular in fault gouges. This unfortunately complicates a straight-
forward application to fault zone problems.
Thermal anomalies associated with seismic slip have been traced during the recent rapid postearthquake
fault drilling projects after major earthquakes, such as the TCDP, the WFSD, and the JFAST projects (respec-
tively, Fulton et al., 2013; Kano et al., 2006; H. Li et al., 2015). In contrast to expectation, only very weak bulk
temperature anomalies (<0.5°C) were identified across these fault zones. The thermal anomalies also appear
to be much lower than predicted from the fault strength deduced from the in situ stress measurements and
laboratory rock friction data (e.g., Brodsky et al., 2020; Fulton et al., 2013; Tanaka et al., 2006; Ujiie
et al., 2013). The temperature within a slip zone is locally elevated by coseismic frictional heating to quite
variable maximum temperatures, even at shallow crustal levels. This is surmised from widely reported ther-
mochemical reactions: clay mineral reactions (e.g., Brantut et al., 2008; L.‐W. Kuo et al., 2011), decarbona-
tion reactions (e.g., Han et al., 2007; Rowe, Fagereng, et al., 2012), amorphization (Aretusini et al., 2019;
Rowe et al., 2019), trace element partitioning features (e.g., Ishikawa et al., 2008; Tanikawa et al., 2015), gra-
phitization of carbonaceous material (e.g., Kuo et al., 2014; Oohashi et al., 2011), thermal maturation of
Figure 13. NRM record in the TCDP fault gouge that hosts the principal slip zone of the 1999 Chi‐Chi earthquake
(Taiwan) (reproduced from Chou, Song, Aubourg, Lee, et al., 2012). (a) Equal‐area stereoplot showing the
Chelungpu fault plane, the mean paleomagnetic components recorded in the three fault zones (FZB1136, FZB1194,
and FZB1243) and the wall rock of the TCDP Hole‐B, and the expected orientation of the earthquake lightning
(EQL) NRM, with an error of ±20° in declination. Solid and open symbols indicate the downward and upward
hemispheres, respectively. The wall rock's main component lies away from the modern magnetic field, which is indicated
as the 1999 international geomagnetic reference field (IGRF) model magnetic vector in central Taiwan by a bold plus
symbol (+). The FZB1136 fault gouge component is closest to the modern magnetic field and statistically different from
the hypothetical EQL direction. In the FZB1194 and FZB1243 gouges, both normal and reversed components
are oriented southerly. (b) Orthogonal plot of AF demagnetization of the NRM in FZB1136 gouge (depth of 1,136.33 m).
Open and solid circles represent projection of the vector on the vertical and horizontal planes, respectively.
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organic molecules (e.g., Coffey et al., 2019; Rabinowitz et al., 2017; Savage et al., 2014), fission track
signatures (e.g., D'Alessio et al., 2003), vitrinite reflectance anomalies (e.g., Maekawa et al., 2014;
Sakaguchi et al., 2011), and occasionally the occurrence of pseudotachylytes (e.g., Otsuki et al., 2009;
Zhang et al., 2017). These diagnostic features provide important constraints on the temperatures in the
PSZ that prevailed during previous earthquake events. However, each of these approaches requires fairly
specific conditions to be applicable. For example, vitrinite reflectance and biomarkers can only be applied
to sediments that contain certain types of organic matter, and fission tracks require the presence of
uranium‐rich inclusions in certain minerals. Also, most approaches only provide information on a certain
temperature threshold or a fairly narrow range in temperature. Vitrinite reflectance, for instance,
indicates temperatures <400°C (Mukhopadhyay, 1992); higher temperatures cannot be diagnosed. Fission
tracks indicate ~120–300°C because they are developed in that temperature window (Figure 14, also see
Yang, Dekkers, & Zhang, 2016, for more details). Thus, to achieve a complete picture of frictional heating,
yet other approaches are desired, complementary to existing methods.
Here, magnetic properties can be used at our advantage. Neoformation of ferrimagnetic minerals through
thermochemical reactions of Fe‐bearing minerals induced by frictional heating is deemed widespread in
seismic fault zones (see section 4.3). This opens up a new perspective in estimating the temperature
experienced during seismic slip (e.g., Chou, Song, Aubourg, Song, et al., 2012; Fukuchi et al., 2005,
2007; Han et al., 2007; Hirono et al., 2009; D. Liu et al., 2016; Mishima et al., 2006; Tanikawa et al., 2007,
2008; Yang et al., 2019, 2018; Yang, Dekkers, & Zhang, 2016). For example, a strong magnetic susceptibil-
ity anomaly in fault gouge was observed at the Fault Zone FZB1136, in TCDP Hole‐B (Figure 11a). It has
been attributed to the formation of large amounts of magnetite (Figures 11c–11f): paramagnetic minerals
(e.g., pyrite, siderite, and chlorite) decomposed at high temperatures (>400°C) during the Chi‐Chi earth-
quake (Chou, Song, Aubourg, et al., 2014; Mishima et al., 2006, 2009). Pyrrhotite was also identified
within the FZB1136 gouge as suggested by FORC diagrams (Figure 15): this provides evidence for thermal
decomposition of pyrite at temperatures >500°C (Chou, Song, Aubourg, Song, et al., 2012). The aforemen-
tioned occurrence of pyrrhotite in slip zones in the Japan Trench frontal prism sediments constrains the
peak temperature of seismic frictional heating to between 640°C and 800°C, combining magnetic infor-
mation with microscopy observations and pyrite‐to‐pyrrhotite reaction kinetics (Yang et al., 2018).
These results provide crucial temperature constraints for understanding the earthquake energy dissipation
and weakening mechanisms, as frictional heating is regarded as the foremost factor controlling fault
weakening (e.g., Bizzarri, 2009; Yao et al., 2016).
Figure 14. Comparison of the proposed provisional rock magnetic “geothermometer”(Yang, Dekkers, & Zhang, 2016)
with other, more conventional, seismic frictional heating thermometers, including vitrinite reflectance, fission‐track,
and biomarkers. The text next to each approach refers to the conditions where the respective thermometers can be
applied. The bar widths indicate their working temperature ranges. The reader is referred to the main text and Yang,
Dekkers, and Zhang (2016) for more details.
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Recently, a preliminary rock magnetic fault zone “geothermometer”was developed, which is utilizing mag-
netic susceptibility versus temperature measurements during cycling in the laboratory to increasingly ele-
vated temperatures (Yang, Dekkers, & Zhang, 2016). The thermochemical expression of magnetic phases
in a rock should not change unless the rock is heated in the laboratory beyond the maximum temperature
it underwent in nature (e.g., Hrouda et al., 2003; Spassov & Hus, 2006). This approach is very similar to esti-
mating temperatures of ancient fires and burnt structures in archeology (Jordanova et al., 2018). In the Japan
Trench subduction plate boundary décollement, cored by the JFAST project, maximum temperatures were
determined ranging from ~300°C to over 500°C close to the multiple slip surfaces of previous earthquakes
(cf. Figure 6b in Yang, Dekkers, & Zhang, 2016). The rock magnetic geothermometer of fault rocks is less
dependent on lithology requirements. Moreover, it offers resolution in the 300–700°C temperature range,
which is difficult to assess with other geothermometers (Figure 14). It thus is a promising complementary
approach for assessing temperature anomalies in slip zones.
Figure 15. (a) Backscatter SEM images of a pyrrhotite‐pyrite grain in the FZB1136 gouge in TCDP Hole‐B, Taiwan. The
framboidal pyrrhotite core has a 1–10 μm pyrite rim. With crossed polars under reflected light microscopy the
pyrrhotite core remains bright and the isotropic pyrite rim is black (insets) (P: cross‐polarized reflected light image;
L: reflected light image with plane polars). Pyrrhotite is inferred to have formed from high‐temperature decomposition of
pyrite (>500°C) during coseismic slip of repeated earthquakes, while pyrite resulted from retrograde metamorphism
of pyrrhotite during cooling of coseismic fluids. (b) FORC diagram for the FZB1136 gouge. A coercivity peak at 90 mT
reveals the presence of magnetically interacting pyrrhotite. (c and d) Modeled Eh‐pH diagrams showing the stability
of pyrite and pyrrhotite at fluid temperatures of 300°C and 100°C inferred for the fault gouge zone of FZB1136 in TCDP
Hole‐B. The pyrite stability field at lower Eh values increases with decreasing temperature from 300°C to 100°C,
whereas that of pyrrhotite field shrinks to a small range with high pH values (8–10) at 100°C, suggesting that pyrite
formed as a result of retrograde alteration of pyrrhotite. The presence of pyrite rims on pyrrhotite cores in the framboids
(Figure 15a) supports this retrograde interpretation. Compiled from Chou, Song, Aubourg, Song, et al. (2012).
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Substantial magnetic enhancement is also observed experimentally in high‐velocity friction experiments on
natural fault gouges with increasing slip distance (e.g., Fukuchi et al., 2005; Tanikawa et al., 2007; Yang
et al., 2019). Magnetite, occurring as spherical and sintered irregular aggregates (Figure 16), was formed dur-
ing friction experiments on fault gouges from the Yingxiu‐Beichuan fault. The original gouges are essentially
paramagnetic with one gouge containing minor goethite (~4 wt%); magnetite is the result of thermochemical
reactions of iron adsorbed on clay minerals (smectite and chlorite) and/or alteration/reduction of goethite.
Higher frictional heating temperatures (equivalent to longer slip distances) led to higher values of magnetic
susceptibility and magnetization and to larger grain sizes of the newly formed magnetite as indicated by
decreasing coercivity (Yang et al., 2019). Changes in magnetic parameters appear to be linear with the
(calculated) temperature rise induced by frictional heating (Figure 16f). Therefore, magnetic properties of
fault rocks can be used not only to detect seismic frictional heating but also to evaluate the amount of
frictional work associated with seismic slip (Tanikawa et al., 2007, 2008; Yang et al., 2019). Clearly, further
experimental work is required to establish a (semi)quantitative relation between magnetic parameters and
the heat production rate (and thus temperature rise) due to sliding. Only then meaningful extrapolation
to natural conditions would be allowed.
However, as discussed in section 4.5, fluid infiltration is widely present in a fault zone. During the postseis-
mic and long‐term interseismic periods, fluids of either meteoric or deep origin (or a combination) may infil-
trate into and percolate through the fault zone. Therefore, the latest coseismically imprinted magnetic
records (i.e., the aforementioned newly formed ferrimagnetic minerals, and consequent high magnetic sus-
ceptibility and magnetization) could be erased or overprinted soon after the earthquake. It is reported that
Figure 16. Rock magnetic properties of experimentally sheared natural fault gouge with a high‐velocity apparatus (compiled from Yang et al., 2019).
(a) A schematic diagram of the rotary shear high‐velocity friction (HVF) testing machine. (b) Sampling strategy after the HVF experiment. The sheared plane is
divided into five annuluses with equivalent slip velocities during the experiments. The slip rate and displacement are 0 at the center and reach maximum
values at the edge. (c) Modeled temperature due to frictional heating in the different annuluses as a function of slip time. The reader is referred to Yang
et al. (2019) for more details on the equations and modeling strategy. (d) Hysteresis loops after and before (insets) correction for the paramagnetic contribution of
the starting material and run products after the experiments. The starting material is dominated by paramagnetic Fe‐bearing minerals; run products contain
an increasing amount of magnetite with slip distance. (e) SEM picture of an Fe oxide spherule with some adhered smaller particles present in the outermost
annulus of the sheared gouge. Inset is EDS spectrum for the analyzed spot marked by the red cross. (f) Magnetic susceptibility of run products linearly increases
with the estimated peak temperatures experienced during the experiment, suggesting enhancement of the newly formed magnetite with increasing
temperature due to frictional heating. GG indicates the starting material; red dots mark the run products. R
2
is the coefficient of determination of the line fits.
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fault properties can be measurably changed 2 years after an earthquake
(Brodsky et al., 2009), and the fluid infiltration would be one of the most
important factors driving these changes (Yang, Yang, et al., 2016). Thus,
there would be a rare opportunity to gain crucial information by observing
such transient magnetic changes related to a specific rupture only imme-
diately after a large earthquake (typically within a couple of years).
5.3. Fingerprint of Fluid Infiltration in Fault Zones
Fluid flow, a major actor in fault zones, is intimately linked to the nuclea-
tion, propagation, arrest, and recurrence of earthquake ruptures (e.g.,
J. Chen et al., 2017; J. P. Evans & Chester, 1995; Goddard &
Evans, 1995; Hickman et al., 1995; Kerrich et al., 1980; Miller, 2013;
Rice, 2006; Terakawa et al., 2010; Williams et al., 2017; Zoback et al., 2007).
Detailed characterization of neoformed and/or altered magnetic minerals
due to fluid‐rock interaction, as discussed in section 4.5 is thus important
for understanding the fluid‐related processes.
In the 16 cm thick gouge of the Fault Zone FZB1136 in TCDP Hole‐B
drilled into the Chelungpu fault in Taiwan, the presence of goethite was
demonstrated by thermal demagnetization of NRM (Figure 11e), mea-
surements of the room temperature SIRM acquired in 2.5 T during cycling
between 300 and 400 K, and transmission X‐ray microscopy (Chou, Song,
Aubourg, Lee, et al., 2012; Chou, Song, Aubourg, Song, et al., 2012). Its
maximum concentration was estimated to be ~1% from the magnetic sus-
ceptibility and SIRM values of the gouge (Figure 11c). The goethite is
argued to have been formed during cooling of a hot coseismic fluid
(>350°C) shortly after the 1999 Chi‐Chi earthquake (Chou, Song,
Aubourg, Lee, et al., 2012; Chou, Song, Aubourg, Song, et al., 2012). The
coseismic fluids were richer in iron in the center of the gouge zone than
at the edges, as demonstrated by the pattern of modeled goethite concen-
tration across the gouge (Figure 11c; see Chou, Song, Aubourg, et al., 2014, for detailed modeling strategy).
On cooling of hot fluids after the earthquake, Eh‐pH conditions are inferred to have changed (Chou, Song,
Aubourg, Song, et al., 2012, Figures 15c and 15d), which led—next to the formation of goethite—to partial
oxidation of magnetite, and some retrograde alteration of pyrrhotite to pyrite (Figure 15a).
Also in the rupture zone of the 2008 Wenchuan earthquake, a marked contrast in magnetic properties was
identified for the two components in the fault breccia, fragments, and matrix. The matrix has a relatively
higher magnetic susceptibility and magnetization but a lower coercivity than the fragments (Figure 17).
The contrast was attributed to fault fluids that caused selective dissolution and precipitation of Fe‐bearing
minerals in the fault breccias during previous earthquake cycles. The matrix is more prone to fluid‐related
effects than the fragments due to its finer grain size and thus larger specific surface (Yang, Chen, et al., 2013).
Magnetite concentrations were found to decrease starting with the host rock, via fault breccia, to (proto)cat-
aclasite. In addition, both low‐and high‐field magnetic susceptibility show close relationships with chlorite
that is a hydrothermal alteration product, as well as with the immobile elements (e.g., TiO
2
and P
2
O
5
; cf.
Figures 13 and 14 in Yang, Yang, et al., 2016). Since those elements are essentially immobile during
fluid‐rock interaction processes, they can be used to infer the losses of mobile elements by mechanical wear
and dissolution, for example, by evaluating patterns in ratios of immobile to mobile elements. Finally, low‐
and high‐field magnetic susceptibility are also positively correlated with the mass losses of fault rocks esti-
mated with the isocon method of geochemical mass balance (e.g., Tanaka et al., 2001). These observations
suggest that magnetite depletion and precipitation of Fe‐bearing clay minerals occurred in these fault
rocks—exhumed from the shallow crust—plumbed by fluid‐assisted processes (Figure 18; see Yang, Yang,
et al., 2016, for details). Hence, magnetic properties of fault rocks represent an expression of fluid infiltration
within fault zones. Further work is necessary to establish a possible causal link between the specifics of mag-
netic minerals and the fault fluid properties (e.g., chemistry, temperature, and pressure).
Figure 17. (a) Sketch of a typical cemented fault breccia from the rupture
zone of the 2008 Wenchuan earthquake (Sichuan Province, China).
The fault breccia is characterized by the formation of host rock derived
fragments embedded within a finer‐grained matrix. (b) A plot of remanent
coercivity (B
cr
) versus the ratio of the saturation remanence and
magnetic susceptibility (M
rs
/χ) shows that the matrix, fragments of the fault
breccias, and bulk samples fall into three groups, suggesting the presence
of different magnetic carriers in the two components (fragments and
matrix) of fault breccias (adapted from Yang, Chen, et al., 2013).
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5.4. AMS and Strain Geometry in Fault Zones
Based on exchange equilibria, various mineral assemblages (e.g., phengite, sphalerite, hornblende, and clin-
opyroxene) have been used as “geobarometer”to define differential stress conditions in crustal rocks (see
Reverdatto et al., 2019, for a review). However, these “stress proxy”minerals are usually absent in fault
zones. This is also the case for more conventional strain markers (e.g., fossils, ooids, pebbles, and nodules)
(Hayman et al., 2004). This makes it difficult to quantify the stress/strain conditions and deformation fabrics
in fault zones. Magnetic fabric analysis, validated for various rock types and in different structural settings
(e.g., Borradaile & Henry, 1997; Hrouda, 1982; Martín‐Hernández et al., 2004; Parés, 2015; Tarling &
Hrouda, 1993), is a widely used tool to examine deformation in geologic structures characterized by large
strain gradients such as shear zones (for a review, see Ferré, Gébelin, et al., 2014). However, to date only a
few studies have targeted brittle fault rocks in the shallow crust (e.g., Casas‐Sainz et al., 2018; Chou, Song,
Lee, et al., 2014; Louis et al., 2008; Marcén et al., 2019; Nakamura & Nagahama, 2001; Román‐Berdiel
et al., 2019; Solum & van der Pluijm, 2009; Yeh et al., 2007). In contrast to the ductile mylonitic rocks, brittle
fault rocks are typified by frictional sliding as dominant deformation process (e.g., Schmid & Handy, 1991;
Sibson, 1977; see section 2 for details). This involves grain rotation, fracturing, and translation of (magnetic)
minerals in fault rocks. The shear effects are reflected in magnetic fabric changes and provide clues on the
strain states in fault zones, as outlined below.
5.4.1. Kinematics of Fault Slip
The orientation of the magnetic fabric provides constraints on the kinematics of fault slip (e.g., Casas‐Sainz
et al., 2018; Elhanati et al., 2020; Ferré et al., 2015, 2016; Hayman et al., 2004; Levi et al., 2014; Marcén
et al., 2019; Solum & van der Pluijm, 2009). In the seismically active Dead Sea Basin (Israel), for example,
the original depositional fabric, characterized by a subvertical κ
min
axis and scattered κ
max
and κ
int
axes, is
well preserved ~2 m away from the fault plane. In contrast, the principal AMS axes of the
deformation‐driven fabrics, which could be observed up to tens of cm from the fault plane, are coaxial with
the instantaneous strain ellipsoids calculated with the fault plane solutions (Figure 19). Such magnetic fab-
rics are interpreted to result from the local strain field during the past earthquakes (Levi et al., 2014).
Another illustrative example is the study of fault rocks from the Chelungpu fault (Taiwan). AMS fabrics
Figure 18. A simple conceptual model illustrating the fluid‐assisted cataclastic process (reprinted from Yang, Yang,
et al., 2016). Cataclasis transforms a granite protolith with magnetite as the dominant magnetic carrier to essentially
paramagnetic fault rocks, that is, cataclasite, fault breccia, and gouge, in the Yingxiu‐Beichuan fault (Longmen Shan
thrust belt, China), which accommodated most of the displacement of the 2008 Mw 7.9 Wenchuan earthquake
(for details the reader is referred to Yang, Yang, et al., 2016).
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revealed that the compressional stress orientation is identical to the regional stress regime and that the mag-
netic foliation planes of fault gouges are consistent with thrust dip‐slip displacement across the gouge zones
(Yeh et al., 2007). Also recently, Ferré et al. (2015) reported for the first time a successful determination of the
full kinematic solution (slip plane, direction, and sense) of a prehistoric (20.1 ± 0.5 Ma) seismic event
through AMS analysis of pseudotachylyte veins in the Dora Maira Massif (Italy; Figure 20). The analysis
relies on the notion that AMS of fault pseudotachylytes arises from coseismic viscous flow of the frictional
melt and therefore is tracking the direction of slip (Ferré et al., 2015). These studies demonstrate that mag-
netic fabrics can be employed not only to detect the local strain field but also to determine fault plane solu-
tions of previous earthquakes.
5.4.2. Strain State Across Fault Zones
Stress and strain information across a fault zone are crucial ingredients for our understanding of earthquake
generation and related slip behavior. Rocks that host (future) faults, in particular unlithified and
Figure 19. Comparison of magnetic fabrics and fault plane solutions across a normal fault in the Lisan Formation,
Masada Plain (Israel), in the Dead Sea basin, where numerous M> 6 earthquake events occurred during the last
70,000 years (compiled from Levi et al., 2014). (a) Lower hemisphere equal‐area projections of principal AMS axes of the
late Pleistocene soft rocks in the footwall (I), intermediate block (II), and hanging wall (III). Dashed lines indicate
the strike direction of the nearby fault strands. The two small stereograms in Figure 19b show the orientations of the fault
strands. (c) Two fault planes and moment‐tensor solutions calculated for the footwall and hanging wall. Dashed
arrows mark the dip‐slip direction. ε
max
,ε
int
, and ε
min
are the instantaneous maximum, intermediate, and least principal
strain axes, respectively. The principal AMS axes of soft rocks close to the three late Pleistocene syndepositional
normal faults resemble the fault plane solutions.
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unconsolidated sediments, respond to the progressive tectonic loading by the development of a preferred
orientation of the constituent (magnetic) minerals. This fabric is commonly coaxial with the axes of the
strain ellipsoid; more specifically, the pole to the magnetic foliation is widely regarded as the orientation
of the maximum shortening strain (e.g., Cifelli et al., 2004; Parés, 2015). Therefore, comparison of the
magnetic fabrics in host rocks bounding the fault zone and fault rocks sheds light on how the strain state
varies across the fault zone. For example, AMS analysis revealed an abrupt change in strain state across
the plate boundary fault zone (décollement) at shallow depths near the Japan Trench (Yang, Mishima,
et al., 2013): The AMS data of the prism sediments above the décollement depict the maximum strain direc-
tion of lateral shortening due to the WNW convergence of the Pacific plate, while those of the underthrust
sediments appear to point to a classic sedimentary compaction fabric, thus a vertical, uniaxial strain
(Figure 21). The plate boundary décollement between the Pacific and North American plates is thus
decoupled over the long term, which favors the propagation of coseismic slip along it (Yang, Mishima,
et al., 2013). Such strain decoupling across a décollement zone revealed by AMS fabrics was also reported
at the Nankai Trough (Owens, 1993; Ujiie et al., 2003), the Costa Rica margin (Housen &
Kanamatsu, 2003), the Barbados accretionary margin (Hounslow, 1990; Housen, Tobin, et al., 1996), and
the Chile Triple Junction region (Collombat et al., 1995). AMS is thus an excellent indicator for incremental
strain that sensitively reflects the stress state across a fault zone.
5.4.3. Deformation Mechanism of Fault Zones in the Shallow Crust
Fault breccia and cataclasite are considered to have been deformed via cataclastic flow (Hayman et al., 2004).
It is a specific deformation mechanism that involves continuous brittle fracturing and comminution
Figure 20. Focal mechanism of a prehistoric earthquake deduced from magnetic fabrics of pseudotachylyte veins
(the Dora Maira massif, western Italian Alps) (compiled from Ferré et al., 2015). (a) Macroscopic view of the
pseudotachylyte generation vein developed in the host mylonitic gneiss. (b) Comparison of magnetic fabrics of the host
gneiss (in black) and pseudotachylyte (in red). The host gneiss (with corrected degree of magnetic anisotropy P
j
of
1.38 ± 0.10) has a subhorizontal magnetic foliation and magnetic lineation. In contrast, the pseudotachylyte foliation
(with P
j
of 1.08 ± 0.01) has an angle of 31° with that of the host gneiss. Itis formed as a result of rapid cooling of the melt
and recorded viscous flow parallel to the seismic slip direction (c), and thus revealing a top‐to‐west sense of shear
and a normal fault focal mechanism.
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of grains with frictional sliding and rolling of the fragments with respect to one another (Twiss &
Moores, 1992). Magnetic grain size reduction occurring during cataclastic flow is expected to decrease the
grain‐scale anisotropy of the AMS carriers (Ferré, Gébelin, et al., 2014; Jackson et al., 1993). This may
explain why notably weak and even nearly isotropic AMS tensors were determined in strongly fractured
granites in the Nojima Fault in Japan (Nakamura & Nagahama, 2001). In contrast, fault gouges are inferred
to be the product of granular flow (Twiss & Moores, 1992) which involves the rolling and sliding of rigid
particles (Morgan, 1999; Morgan & Boettcher, 1999). In this process, the physical rearrangement of the mag-
netic minerals primarily results in an increase in the AMS parameters (e.g., Pand F). One should bear in
mind, however, that brittle fault rocks often host secondary minerals that not necessarily relate directly to
strain. For example, in the Fault Zone FZB1136 in TCDP Hole‐B the magnetic foliation and anisotropy
degree are lowest in the PSZ itself, whereas the highest values appear in the fault gouge where goethite is
present (Chou, Song, Lee, et al., 2014). The goethite was formed postseismically by the action of
thermal fluids (Chou, Song, Aubourg, Song, et al., 2012, see also section 5.3). Therefore, caution is appropri-
ate when interpreting the deformation mechanisms based on the AMS analysis of brittle fault rocks (e.g.,
Román‐Berdiel et al., 2019). Microscopic inspection of whether or not secondary minerals are present is
recommended.
Figure 21. (a) The general plate configuration of the Japanese island arc and map of the 2011 Mw 9.0 Tohoku‐oki earthquake portion of the Japan Trench,
showing the location of the IODP Expedition 343 drill site with a filled white circle. The red star indicates the epicenter of the earthquake. White arrow indicates
the Pacific plate (PAC) convergence vector with a rate of∼85 mm/year (Argus et al., 2011). (b) Downcore profiles of degree of anisotropy (P) and magnetic
foliation (F). Both progressively increase with depth in the frontal prism and suddenly drop just below the décollement, mostly showing small values in the
underthrust sediments. Equal‐area, lower hemisphere projections of the bootstrapped κ
max
,κ
int
, and κ
min
directions for (c) the prism sediments (n= 34) and (d)
the underthrust sediments (n= 22). The AMS fabric pattern of the prism sediments is consistent with horizontal tectonic shortening (red arrows in Figure 21c)
nearly parallel to the PAC convergence direction in the Japan Trench, as well as the maximum horizontal stress deduced from borehole breakouts
(W. Lin et al., 2013). The AMS in the underthrust sediments represents a vertical, uniaxial strain, that is, an initial sedimentary compaction fabric. These very
different AMS patterns reveal an abrupt strain change across the plate boundary décollement at shallow depths and imply that large coseismic slip occurred along
a weak décollement that is mainly decoupled over the long term. Compiled from Yang, Mishima, et al. (2013) with data from Argus et al. (2011) and W. Lin
et al. (2013).
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The magnetic fabric of pseudotachylytes is providing a snapshot of the ambient stress during or immediately
after seismic slip (e.g., Ferré et al., 2015, 2016; Molina‐Garza et al., 2009). In the Chiapas Massif (Mexico), for
example, the magnetic fabrics of the host granitoid rock and the pseudotachylyte veins are essentially
parallel, with nearly east‐west and steep magnetic foliations. However, the AMS fabric of the host rock is
dominantly prolate, while that of the pseudotachylytes veins is markedly oblate. This suggests that the veins
underwent a pure shear deformation (i.e., viscous breaking) during cooling (Molina‐Garza et al., 2009),
again illustrating that AMS analysis delivers important insights into the deformation mechanism of
fault zones.
5.5. Magnetic Records of Coseismic Electric Currents
As mentioned earlier, several faulting‐related magnetization processes contribute to the NRM acquisition of
pseudotachylytes, including coseismic TRM acquired upon cooling of the melt below the Curie temperature,
EQL‐induced IRM acquisition (EQLIRM), and postseismic CRM due to both devitrification and alteration
(Ferré et al., 2012; Leibovitz, 2016). Deciphering these magnetization components is crucial to the under-
standing the pseudotachylytes' microstructures and the timing of their development.
Coseismic TRM is restricted to the local melt pockets and records the direction and intensity of the geomag-
netic field at the time of cooling. It is straightforwardly demagnetized thermally in the laboratory. In con-
trast, EQLIRM typically displays an anomalously high NRM value in pseudotachylytes with characteristic
directions at a high angle to the fault plane. There is no geometric relationship to the past or present direc-
tions of the geomagnetic field (Leibovitz, 2016). Thus, rock magnetic and paleomagnetic analyses of the
pseudotachylytes do provide independent constraints on the origin(s) of the NRM in pseudotachylytes.
Anomalously high NRM recorded in a pseudotachylyte, which would require a magnetic field (much) stron-
ger than the typical Earth's magnetic field, then points to EQLIRM as the dominant NRM acquisition
process.
Indeed, such abnormally high NRM intensities have been reported for fault rocks (fault gouge and pseudo-
tachylytes) along or near the fault plane where large earthquakes occurred (e.g., Enomoto & Zheng, 1998;
Enomoto et al., 2001; Ferré et al., 2005, 2012; Ferré, Geissman, et al., 2014). For instance, pseudotachylyte
with an abnormally high NRM (~132 A/m), up to 300 times higher than that of the granitic host rock
(tonalite), was found in the Santa Rosa Mountains near Palm Springs, California, USA (Ferré et al., 2005).
These pseudotachylyte veins typically have a moderate‐coercivity, well‐defined, single component NRM iso-
lated with stepwise progressive alternating field demagnetization (cf. Figure 6a in Ferré et al., 2005). For
comparison, typical NRM values of volcanic rocks that cool in the Earth's magnetic field range from 0.1 to
10 A/m depending on their chemical composition. Mafic and ultramafic rocks are more magnetic than felsic
rocks. To acquire the reported high NRM, the fault rocks and pseudotachylytes must have been exposed to a
strong local magnetic field, several orders of magnitude higher than the typical Earth's main field (Ferré
et al., 2005). The most likely source of the required strong magnetic field is believed to be the coseismic elec-
tric currents produced by EQL (Ferré et al., 2005). In the Nojima Fault (Japan), the coseismic electric current
was estimated to be as high as ~1 kA during the 1995 Kobe earthquake through simulation of the
EQL‐induced sintering of the fault gouge using spark plasma sintering experiments (Enomoto et al., 2001).
Such a large current is expected to produce a strong pulsed magnetic field, in which the foliated fault gouges
acquired an IRM which resulted in a very large NRM: ~430 times higher than that of the granitic fault brec-
cia, and ~70 times higher than that of the adjacent mudstone (Enomoto et al., 2001; Enomoto &
Zheng, 1998).
6. Comparison of Magnetic Methods to Other Approaches
The magnetic properties of fault rocks can provide intriguing insights into faulting processes. The number of
magnetic studies on fault rocks, however, is rather limited to date. A broad‐scale comparison of fault rock
magnetic data with other more classical approaches is therefore currently in its infancy. Below we assemble
the information concerning the temperature of frictional heating and the stress/strain state of the fault
retrieved from magnetic and nonmagnetic thermal and strain indicators. This comparison is only possible
in a few fault zones that accommodated large earthquakes, where data from multiple techniques are
available.
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6.1. The Seismic Frictional Heating Temperature
Frictional heating during seismic slip is evidently registered in all minerals in the slip zone, not only in mag-
netic minerals. Nonmagnetic minerals include hydrous, clay, and carbonate minerals. Thus, frictional heat-
ing might be traced and cross calibrated by a variety of thermal indicators (Table 1). In the Fault Zone
FZ1136, TCDP Hole‐B, the peak temperature during the 1999 Chi‐Chi earthquake was >400–500°C, inferred
from the thermal decomposition of paramagnetic minerals with the formation of ferrimagnetic minerals
(Table 1; Chou, Song, Aubourg, Song, et al., 2012; Mishima et al., 2006, 2009; Tanikawa et al., 2007, 2008).
This estimate concurs with others based on nonmagnetic thermal indicators: The signature of mobile trace
elements in the fault fluid indicates a frictional heating peak temperature >350°C (Ishikawa et al., 2008).
That temperature is estimated at >300–400°C based on the vaporization of water during thermal pressuriza-
tion (Boullier et al., 2009). These peak temperatures estimates are lower than the estimate of 626 ± 25°C
from a kinetic model of vitrinite thermal maturation (Maekawa et al., 2014), and the infrared and Raman
estimates from carbonaceous fault materials (≥700°C, Hirono et al., 2015). The clay mineral signature would
indicate that the peak temperature could even be up to 900–1100°C (L.‐W. Kuo et al., 2011). Most recently,
nanometric geochemical analysis of fault gouge revealed that most of the siderite in the slip zone has
been evaporated into nanosized grains, suggesting that the peak temperature reached as high as ~1200°C
(W.‐H. Li et al., 2019). While the actual peak temperature is still not well defined, the existence of frictional
heating has been definitively verified by the 0.06°C residual temperature anomaly measured 6 years after the
1999 Chi‐Chi earthquake rupture through direct borehole temperature logging (Table 1; Kano et al., 2006). A
similar situation exists in the Japan Trench subduction plate boundary fault zone which accommodated the
large slip of the 2011 Tohoku‐oki earthquake (Table 1) with a thermal of anomaly of 0.31°C measured
16 months after rupture (Fulton et al., 2013).
Table 1
Comparison of Peak Temperatures of Frictional Heating Estimated by Magnetic and Nonmagnetic Approaches in Three Fault Zones That Accommodated Large
Magnitude Earthquakes
Peak temperatures Indicators/methods Note
TCDP Hole B Fault Zone FZ1136; hosted the 1999 Chi‐Chi (Taiwan) earthquake (Mw 7.6)
>400°C Thermal decomposition of paramagnetic minerals
a,b,c,d
Magnetic indicators
>500°C Pyrrhotite formation via thermal breakdown of pyrite
e
>350°C Fluid‐mobile trace element anomalies
f
Nonmagnetic indicators/methods
>300–400°C Vaporization of water during thermal pressurization
g
626 ± 25°C Vitrinite reflectance
h
~≥700°C Infrared and Raman spectra of carbonaceous materials
i
900–1100°C Clay mineral anomalies (FZ1111 in Hole A that corresponds to FZ1136 Hole B)
j
~1200°C Thermal decomposition or breakdown of siderite into nanosized grains
k
0.06°C residual temperature anomaly
6 years after the earthquake
Direct borehole observation
l
Japan Trench subduction plate boundary fault zone; hosted the 2011 Mw 9.0 Tohoku‐oki earthquake
>300–500°C Rock magnetic “geothermometer”
m
Magnetic indicators
640–800°C Reaction of pyrite to pyrrhotite
n
120–900°C Organic biomarker thermal maturity
o
Nonmagnetic indicators/methods
0.31°C residual temperature anomaly
16 months after the earthquake
Direct borehole observation
p
<1250°C Modeled with plausible slip durations and slip zone thicknesses, and a
friction coefficient of 0.08
p
Nankai Subduction Zone (Japan) associated with the 1944 Tonankai Mw 8.1 earthquake
<400°C No thermal decomposition of paramagnetic minerals occurred
q
Magnetic indicators
<300°C Fluid‐mobile trace elements, Raman spectra of carbonaceous material,
and inorganic carbon content
q
Nonmagnetic indicators/methods
390 ± 50°C Vitrinite reflectance
r
<250°C Infrared spectroscopy, clay mineral anomalies, trace‐element geochemistry,
and isotope geochemistry
s
a
Mishima et al. (2006).
b
Mishima et al. (2009).
c
Tanikawa et al. (2007).
d
Tanikawa et al. (2008).
e
Chou, Song, Aubourg, Song, et al. (2012).
f
Ishikawa
et al. (2008).
g
Boullier et al. (2009).
h
Maekawa et al. (2014).
i
Hirono et al. (2015).
j
L.‐W. Kuo et al. (2011).
k
W.‐H. Li et al. (2019).
l
Kano et al. (2006).
m
Yang, Dekkers, and Zhang (2016).
n
Yang et al. (2018).
o
Rabinowitz et al. (2020).
p
Fulton et al. (2013).
q
Hirono et al. (2009).
r
Sakaguchi et al. (2011).
s
Hirono et al. (2014).
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The peak temperature estimates obtained by different methods thus appear to be quite variable. A possible
reason could be that different thermal proxies sense correctly certain temperature intervals only
(cf. Figure 14, see also Rowe & Griffith, 2015; Yang, Dekkers, & Zhang, 2016). Another reason could be that
the majority of the temperature estimates take equilibrium conditions as their basic premise, which may not
apply to the timescales of fault slip (e.g., Savage et al., 2018). This clearly warrants further work. However, at
present we may deduce that peak temperature estimates of fault rocks inferred from magnetic properties are
reasonably in line with the majority of other methods and provide at least a lower bound on the temperature
rise due to seismic frictional heating.
In other cases examination of magnetic assemblages can provide an upper limit for the temperature of the
frictional heating. This is the case in the Nankai subduction zone (Japan) associated with the 1944
Tonankai Mw 8.1 earthquake, where the absence of thermal decomposition of paramagnetic minerals indi-
cates a peak temperature <400°C (Table 1; Hirono et al., 2009). This peak temperature estimate is corrobo-
rated by estimates from other thermal indicators, such as the fluid‐mobile trace element signature, the
Raman spectral features of carbonaceous material, and the inorganic carbon content (these parameters all
indicate <300°C, Hirono et al., 2009, 2014). Also, vitrinite reflectance data (peak temperature 390°C,
cf. Sakaguchi et al., 2011) support the magnetic information. It further shows that magnetic property
analysis of fault rocks can provide at least a reasonably robust thermal record of earthquake slip. It should
be realized that the amount of frictional heating can vary widely for different earthquakes; the related
temperature rise is as yet rather poorly constrained by any individual method. Therefore, all approaches
are subject to ongoing testing and refinement. Combined application, reconciliation and cross calibration
of different thermometers would thus yield the most robust picture. This is also where the magnetic methods
will be playing an important role.
6.2. Stress Orientations
In the Japan Trench subduction fault zone, the maximum AMS axes of sedimentary rocks in the frontal
prism align northeast‐southwest (36.6 ± 11.3°, Figure 21c, Yang, Mishima, et al., 2013). This is nearly per-
pendicular to the direction of the Pacific plate convergence (~292°, Argus et al., 2011) and the maximum hor-
izontal stress orientation (319 ± 23°) deduced from the azimuths of borehole breakouts (W. Lin et al., 2013;
see Appendix A for definition). The accretionary prism sediments thus probably have undergone layer par-
allel shortening due to the convergence of the Pacific plate. Similar AMS patterns have also been reported in
the TCDP core samples, where the magnetic ellipsoids are coaxial with the (palaeo)stress ellipsoids deduced
from fold axes, fault analysis, and borehole breakouts (cf. Figure 14b in Louis et al., 2008). Hence, these
observations provide a strong validation of the reliability of AMS as a (palaeo)stress indicator within fault
systems.
7. Remaining Challenges
The foregoing discussion illustrates that the rock magnetic approach is very promising to unveil valuable
information pertaining to faulting processes. Comparison of magnetic and nonmagnetic approaches, as
low as the number of comparative studies at this moment, yields a broad agreement among most of the
approaches. Moreover, they have improved our understanding of the magnetic signatures of fault rocks.
Nonetheless, some essential questions pertaining to the mechanisms that underlie the magnetic changes in
fault rocks remain currently unanswered. Most importantly, the reaction kinetics of Fe‐bearing (magnetic)
mineral reactions under the extreme faulting conditions (i.e., fast heating, intensive mechanical milling, and
nanocrystallization) are not well understood. On the other hand, other than the protolith lithology, magnetic
properties of fault rocks are strongly affected not only by physicochemical conditions (temperature, pres-
sure, oxygen fugacity, pH, etc.) during rupture but also by alteration and chemical differentiation within a
fault zone during the other stages of the seismic cycle, that is, the postseismic, and interseismic periods
(e.g., Bense et al., 2013; Ferré et al., 2017; O'Hara & Huggins, 2005; Yamaguchi et al., 2011). This complicates
the magnetic signature interpretation of fault rocks. The individual impact of each stage of the seismic cycle
must be untangled before a full understanding of faulting behavior is obtained. For example, fluid‐rock
interaction plays a key role in the alteration, removal, and/or neoformation of Fe‐bearing minerals in fault
zones. Unfortunately, the associated kinetics are presently rather poorly understood (e.g., Chou, Song,
Aubourg, Song, et al., 2012; Yang, Chen, et al., 2013; Yang, Yang, et al., 2016). Also, identification of a
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PSZ and the signature of frictional heating of individual earthquake events are not “one to one”, as a fault
zone is the cumulative product of many earthquake cycles, and thus, several PSZs may stack together in a
fault gouge. These challenges warrant future work.
Next, we identify some attractive research avenues to help addressing these concerns, including (1) extensive
magnetic analysis on fault rocks after friction experiments, (2) laboratory simulation of fault fluid percola-
tion, (3) emerging technology for resolving faulting imprinted NRM components, (4) magnetic fabric analy-
sis of brittle fault rocks, and (5) synergy of interdisciplinary approaches in mineral‐magnetic studies.
7.1. Extensive Magnetic Analysis on Fault Rocks After Frictional Experiments
Friction impacts the formation and/or destruction of Fe‐bearing (magnetic) minerals. Mechanical milling in
air can completely transform hematite to magnetite (Petrovský et al., 1996, 2000; Zdujićet al., 1998);
mechanical friction in argon favors the thermal decomposition of pyrite into pyrrhotite (H.‐P. Hu et al., 2002)
while mechanochemical decomposition of siderite due to crushing may produce magnetite (Criado
et al., 1988). Similar processes can readily occur in seismic slip zones during earthquakes (e.g., J. P. Evans
et al., 2014). Experimental calibration of the reaction kinetics on time‐temperature paths similar to those
expected on fault surfaces during seismic slip may enable precise determination of the concomitant tempera-
ture rise. Therefore, laboratory high‐speed friction experiments using rotary shear apparatus (e.g., Di Toro
et al., 2010; S. Ma et al., 2014) on natural and synthetic fault gouges are promising to precisely constrain
the processes that occurred in a fault zone during rupture. Such experiments could be carried out on
samples with various compositions focusing on the suite of Fe‐bearing minerals and under different condi-
tions (e.g., variable oxygen fugacity, permeable/impermeable, high vapor pressure, hydrostatic, and elevated
temperature). Real‐time records of physical parameters, that is, temperature, velocity, stress, displacement,
provide high‐resolution information on the “flash‐heating”effects due to the friction. Therefore, studies
should ideally integrate mineral magnetic, mineralogical, geochemical, and microscopic aspects on fault
gouges before and after experimental shearing. Also, they should be complemented by kinetic chemical
modeling and temperature modeling. With such data, our understanding of the evolution of Fe‐bearing
(magnetic) mineral(s) within the PSZ during slip will be put on a much stronger footing. In addition, similar
experiments under low‐speed conditions (as low as the mm/year range, cf. Collettini et al., 2014, and S. Ma
et al., 2014) may shed light on the possible magnetic response during earthquake nucleation.
7.2. Laboratory Simulation of Fluid Percolation
Fluids along fault zones are generally of multiple sources (Zoback et al., 2007). Their migration paths and
compositions vary markedly over time and space, as fault zones act as fluid conduits or barriers at different
stages in their development, that is, the coseismic, postseismic, and interseismic periods (Bense et al., 2013).
This leads to temporal and spatial changes of the redox state in a fault zone as a function of episodic fault
rupture and healing cycles during the lifetime of a fault (Yamaguchi et al., 2011). Laboratory simulation
of such processes with natural and/or synthetic fault rocks under controlled physicochemical conditions
(i.e., temperature, pressure, redox state, pH, and fluid‐rock ratio) thus would be a promising way to improve
our understanding of alteration, removal, and/or neoformation of magnetic minerals as a result of fluid infil-
tration. Answers may be retrieved from characterizing the Fe‐bearing and associated minerals which are
precipitated or recrystallized, such as iron oxides, iron hydroxides, iron carbonates, and/or iron sulfides,
in incorporation with thermochemical equilibrium modeling under the specific experimental physicochem-
ical conditions.
7.3. Emerging Technology for Resolving Faulting Imprinted NRM Components
Isolation and characterization of faulting‐induced NRM components are expected to yield valuable informa-
tion on faulting‐related effects. With the exception of a few early studies on paleomagnetic dating of fault
rocks (e.g., Grønlie & Torsvik, 1989; Hailwood et al., 1992; Torsvik et al., 1992); however, little attention
has been paid to newly imprinted “secondary”coseismic/postseismic NRM components, such as related
(p)TRM or CRM (e.g., Chou, Song, Aubourg, Lee, et al., 2012, see section 5.3) and potential EQLIRM
(e.g., Ferré et al., 2012; see section 5.5). This is most likely dictated by the amount of sample material needed
for the required precise determination of paleomagnetic signatures by classic paleomagnetic measurements.
Alternative options include the aforementioned SSM and QDM. The SSM is capable of mapping magnetic
fields with high spatial resolution (better than 100 μm at room temperature and 4 μm or better at
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cryogenic temperatures), with a limiting magnetic moment sensitivity of ~1 × 10
−14
Am
2
at a sensor‐to‐sam-
ple distance of 200 μm (e.g., Fukuzawa et al., 2017; Pastore et al., 2018). It has been successfully applied in
various paleomagnetic studies, such as on lunar glass spherules and Hawaiian basalt (Weiss et al., 2007) and
ferromanganese crust (Noguchi et al., 2017). In the fault slip context, magnetic stripes with positive and
negative polarities subparallel to the slip surface were recently detected by Fukuzawa et al. (2017) on the
Nojima Fault gouge (Japan) within an area of 1 cm × 1 cm, using SSM. Such NRM distribution patterns
in fault gouge could provide a means to distinguish prehistorical periods of earthquake activity on the slip
zone with the help of the geomagnetic polarity timescale. The QDM is a combination of superior spatial reso-
lution (5 μm), magnetic sensitivity (20 μTμm/Hz
1/2
), and wide field of view (4 mm) (Glenn et al., 2017). It
can image both remanent and induced magnetization of geological samples at room temperature (Glenn
et al., 2017). These emerging microscopic magnetic scanning techniques enable micrometer‐scale paleomag-
netic studies of fault zones. They also offer spatially well‐resolved measurements of magnetic signals of the
narrow fault zones. Thus, they represent a very promising way to resolve the cumulative frictional heating
effects in PSZs caused by recurrent (historical) earthquakes.
7.4. Magnetic Fabric Analysis of Brittle Fault Rocks
Fault zones are most complex study targets. The caveats put forward by Ferré, Gébelin, et al. (2014) for duc-
tile shear zones also apply to brittle fault rocks in the shallow crust. First, mineral assemblages and mechan-
ical properties of fault zones are heterogeneous between and within the different fault zone compartments
that is, damage zone, fault breccia, and fault gouge. The AMS fabrics in each zone could well have their indi-
vidual strain relationship. Meanwhile, faulting‐related process (frictional heating and fluid‐rock interac-
tions) tends to induce alteration of mineral assemblages that would not necessarily relate to strain (e.g.,
Just et al., 2004). The observed changes in AMS thus cannot be solely attributed to deformation and evidently
complicate the relation of AMS with strain in a fault zone. As an example, the AMS fabric of fault gouges in
the Nojima Fault (Japan) is isotropic (Nakamura & Nagahama, 2001). However, it is anisotropic in the
Chelungpu Fault in Taiwan (Yeh et al., 2007). Integrated rock magnetism, mineralogy and geochemistry,
microstructural observations, and possibly the determination of other magnetic fabrics (high field and low
temperature) and the AMR analysis are desired to fully unravel the relation between magnetic fabric and
the deformation mechanisms operating during faulting.
In addition, the neoformed magnetic minerals and paramagnetic/diamagnetic minerals in fault rocks may
well be formed at different faulting stages or events and could have distinct deformation paths. For example,
diamagnetic fabric of coccolith calcite from the plate boundary Dead Sea Fault indicates tectonic strain,
while the paramagnetic fabric of clays preserves the depositional fabric (Issachar et al., 2018). Thus, evalua-
tion of the relative contributions from the ferrimagnetic, paramagnetic, and diamagnetic components to the
bulk AMS (e.g., Ferré et al., 2004; Hrouda & Jelínek, 1990; Kelso et al., 2002; Martín‐Hernández &
Ferré, 2007) and examination of the (weak) field dependence of AMS, as well as the out‐of‐phase and in‐
phase components of AMS, could be insightful into the dynamic deformation processes in a fault zone.
7.5. Synergy of Interdisciplinary Approaches in Mineral‐Magnetic Studies
The various structural domains of a fault zone that is, fault core, damage zone, and host rock, have a very
diverse level of faulting effects and number of rupture episodes. Attention should thus be paid not only to
bulk magnetic properties of these different fault rocks (e.g., fault breccia, cataclasite, and fault gouge) but
also the different components of fault rocks should be considered individually. Fault breccia, for instance,
generally comprises a mixture of fragments with different grain sizes, set in a finer‐grained, clay‐rich matrix
(Figure 17a). With their heterogeneous hydraulic structure and specific surface, alteration associated with
faulting processes (fluid percolation in particular) acts on them to quite different extents (Yang,
Chen, et al., 2013; Yang, Yang, et al., 2016). Consequently, an assessment of the relative contributions of dif-
ferent components in fault rocks to the overall magnetization, and unraveling their individual alteration
processes are of great help for understanding the changes in magnetic minerals as a fault zone evolves.
This could be achieved through combination with mineralogical, geochemical, and microscopic analyses,
other than mere rock magnetic examination of the iron‐bearing minerals (e.g., Fe oxides, Fe hydroxides,
Fe sulfides, Fe carbonates, or Fe‐silicates) in fault materials.
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The thickness of PSZ and microstructures in fault zones is often at the millimeter to even micrometer scale
(e.g., Boullier, 2011; Fondriest et al., 2013; Sibson, 2003; Siman‐Tov et al., 2013). Thus, with magnetic studies
at a microscale one should be able to gather most meaningful information on temperature changes within a
PSZ. Next to the above‐mentioned SSM and QDM, advanced techniques for magnetic imaging at microscale/
nanoscale are recommended in future studies, including but not limited to magnetic tunnel junction for
mapping magnetic fields with high spatial resolution and field sensitivity (e.g., Lima et al., 2014) and
synchrotron radiation analysis that enables imaging the three‐dimensional structures of nanosized magnetic
minerals (e.g., Chou, Song, Tsao, et al., 2014), as well as the deformation microfabrics in fault zones
(e.g., Fusseis et al., 2014). Orientation statistics from three‐dimensional structure of magnetic particles then
become valuable, which could be matched with, for example, AMS or other magnetic techniques (e.g., Petri
et al., 2020).
Figure 22. Aflowchart for the magnetic property analysis of fault rocks. Multidisciplinary and integrated approaches
incorporating field studies, laboratory simulations, high‐resolution microscopy, and numerical modeling are proposed as
a most promising way to gain a full appreciation of the magnetic response to dynamic physicochemical processes in
fault zones. With a research effort along such lines, better rock magnetic “strain indicators,”“geothermometers,”and
“fluid tracers”are expected to become available for fault zone studies. XRD, X‐ray diffraction; XRF, X‐ray fluorescence;
SEM, scanning electron microscopy; TEM, transmission electron microscopy; EDS, energy dispersive spectroscopy;
EPMA, electron probe microanalysis; ESR, electron spin resonance; DRS, diffuse reflectance spectroscopy; MFM,
magnetic force microscopy; SSM, scanning SQUID microscopy; and QDM, quantum diamond magnetometry.
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8. Conclusions
We have examined the pros and cons of rock magnetic methods applied to fault zones. After identifying can-
didate mechanisms for magnetic changes in fault rocks, we summarize the recent advances of such studies
on fault rocks, highlighting the insights that were gained from these studies. It has been demonstrated that
the rock magnetic properties can be employed as “strain indicator”,“geothermometer”, and “fluid tracer”in
fault zones. This opens a new window into investigating faulting processes.
This emerging subdiscipline is promising and challenging at the same time. Magnetic properties of fault
rocks hold many secrets that await discovery; multidisciplinary and integrated approaches are crucial to gain
a full appreciation of magnetic response to dynamic physicochemical processes in fault zones (Figure 22).
Other than magnetic investigations of fault rocks from earthquake faults, laboratory experiments (friction
experiments, milling, rapid thermal treatment, and fluid percolation, etc.) on both natural and synthetic
fault rocks should be taken into consideration. Additionally, kinetic, equilibrium, and thermodynamic mod-
eling incorporating the results of field studies and laboratory simulations should be included. All of these
efforts would serve to unraveling the mechanisms responsible for magnetic changes in a fault zone, and con-
tribute to the development and refinement of rock magnetic “strain indicators”,“geothermometers”, and
“fluid tracers”in fault zones. This will ultimately yield a better tool for illuminating faulting processes.
Appendix A: Glossary
A.1 Part I: Rock Magnetic Terms
•Anisotropy of (low‐field) magnetic susceptibility (AMS): The low‐field susceptibility varies slightly as a
function of orientation of the sample with respect to the applied field (with a strength of up to several
times the Earth's magnetic field). This variation is described by a symmetric tensor, the AMS tensor.
AMS reflects the statistical alignment of crystallographic directions and/or shape in platy or elongate
grains. Next to AMS also the anisotropy of various remanences imparted in the laboratory is occasionally
evaluated, mostly the AARM.
•Antiferromagnetic: A state of magnetic order with two ferromagnetic sublattices of equal size aligned anti-
parallel so that (ideally) no net magnetic moment results. In practice, a small magnetic moment persists
because of small deviations from the perfect antiparallel state or defects in the crystal structure. Example:
hematite.
•Chemical remanent magnetization (CRM): The NRM imparted to magnetic minerals by chemical pro-
cesses, at any temperature below their Curie temperatures, in the presence of an ambient geomagnetic
field. Chemical processes involve grain growth beyond the SP grain size or thermochemical alteration
of existing magnetic minerals to other magnetic minerals.
•Coercive force (B
c
) and coercivity of remanence (B
cr
): see magnetic hysteresis loop and parameters.
•Curie temperature (T
C
): As the temperature increases in a ferromagnetic (sensu lato) material, interatomic
distances increase, and the magnetic exchange interaction that describes the collective ordering of atomic
magnetic moments becomes weaker. At T
C
, the thermal energy overcomes the exchange energy, and indi-
vidual atomic magnetic moments become independent (they lose their long‐range ordering) so that the
material becomes paramagnetic. Named after the French physicist Pierre Curie (1859–1906).
•Diamagnetic: A material is diamagnetic when all electron spins are paired. When exposed to a magnetic
field, a diamagnetic material tends to reduce that magnetic field: Its susceptibility is negative (the induced
magnetization is opposite to the applied field). The value is very small and independent of temperature.
Examples: quartz (SiO
2
) and calcite (CaCO
3
).
•Ferrimagnetic: A state of magnetic order in which two ferromagnetic sublattices of unequal magnitude are
aligned antiparallel to each other. Example: magnetite.
•Ferromagnetic: A state of magnetic order in which all atomic magnetic moments are aligned parallel to
each other. Example: metallic iron.
•First‐order reversal curve (FORC) diagrams: FORCs are partial hysteresis loops determined with the fol-
lowing procedure: (1) a positive saturation field is applied, (2) the field is ramped down to a lower reversal
field, and (3) the magnetization of the sample is measured, while the applied field is ramped upward to
positive saturation again. A series of FORCs is measured for a set of reversal fields that are increasingly
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away from positive saturation, and a FORC distribution is calculated as the mixed second derivative of
magnetization with respect to field spacing; it is displayed in a so‐called FORC diagram. FORC diagrams
provide information about the distribution of switching fields (i.e., coercivities) and interaction fields for
all magnetic particles that contribute to a hysteresis loop. They are widely used to characterize samples in
rock and mineral magnetism.
•Frequency‐dependent magnetic susceptibility: The difference between susceptibility measurements made
at two frequencies (e.g., 0.47 and 4.7 kHz), expressed either as a percentage of the low frequency or as
an absolute susceptibility value (i.e., in m
3
kg
−1
for mass‐specific susceptibility or in dimensionless
units for volume‐specific susceptibility). This measurement denotes the presence of magnetically “vis-
cous”ferrimagnetic (magnetite or maghemite) grains lying at the stable single domain (SSD)/SP
boundary.
•Hysteresis parameters: see magnetic hysteresis loop and parameters.
•In‐phase and out‐of‐phase magnetic susceptibility: When measuring susceptibility in low alternating mag-
netic fields, the measured specimen is magnetized by a weak field sinusoidally varying in time: H
(t)=H
0
cos(ωt), where H
0
is the field amplitude, ωthe angular frequency, and tthe time. The magnetic
response is measured, most conveniently represented by magnetization M(t). In diamagnetic, paramag-
netic, and many ferromagnetic (sensu lato) materials, the magnetization also varies sinusoidally being
in phase with the applied field. The magnetization to field ratio then maintains a constant value that
represents the susceptibility. In some ferromagnetic materials, however, the magnetization is not in phase
with the applied field but lags behind the field usually because of electrical conduction. In that case M′
and M″are the in‐phase and out‐of‐phase magnetization components, respectively. Then, the in‐phase
susceptibility is χ′=M′/H
0
and the out‐of‐phase susceptibility is χ″=M″/H
0
. See Hrouda et al. (2016)
for more details.
•Magnetic domain: A portion of a ferromagnetic/ferrimagnetic material in which the atomic magnetic
moments are aligned. In a magnetic material, magnetic domains serve to reduce the magnetostatic energy
due to the aligned spontaneous magnetization. Fine magnetic grains are SD, and larger grains contain
several magnetic domains. The magnetic domains form a structure in which the magnetic moments of
individual magnetic domains cancel each other out as much as possible. As a result, the spontaneous mag-
netization of the overall magnetic body is reduced, leading to a lower magnetostatic energy state which
represents a most stable energy configuration.
•Magnetic susceptibility: A measure of the ease with which a substance can be magnetized in small, Earth‐
like, applied magnetic fields. The magnetic moment divided by the field strength in units of H(i.e., A m
−1
)
is the susceptibility. Volume susceptibility (κ) is the induced magnetization per unit volume divided by the
field strength, and is a dimensionless quantity. Mass‐specific susceptibility (χ) is the magnetic moment
expressed per unit of mass divided by the field strength; it has units of m
3
kg
−1
, that is the inverse of
the specific density of a material. Magnetic susceptibility depends principally on the type and concentra-
tion of magnetic minerals in a sample; grain‐size dependence is marginal (with the exception of superpar-
amagnetism that occurs below ~25 nm at room temperature).
•Magnetic remanence: Magnetization that persists in the absence of an applied magnetic field. It differs
from induced magnetization that disappears on removal of an applied field.
•Magnetic hysteresis loop and parameters: A hysteresis loop results when the in‐field magnetization of a
sample is measured as the inducing field is cycled between high positive and negative values. A hysteresis
loop can be measured at any temperature; most hysteresis loops are acquired at room temperature. The
shape of the hysteresis loop yields pertinent information about the magnetic mineral(s) present in a sam-
ple and most importantly on their grain size. The maximum‐induced magnetization that can be induced
in a sample is termed the saturation magnetization (M
s
). The corresponding remanence, when the applied
field is reduced to 0, is the saturation remanence (M
rs
), and the reversed magnetic field required to reduce
the magnetization to zero is the coercive force (B
c
). The coercivity of remanence (B
cr
): The oppositely
directed magnetic field required to demagnetize the saturation remanent magnetization to zero. The
M
rs
/M
s
ratio is often plot against the B
cr
/B
c
ratio, named the Day plot (after Ron Day) which features mag-
netic grain size regions. Without additional information on the magnetic mineralogy and their oxidation
state interpretations in terms of grain size should be exercised with caution.
•Magnetic saturation: The maximum magnetic moment a sample can achieve. Increasing the field strength
even further does not impart a higher magnetic moment.
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•Magnetic vortex state: The transitional state between SD and MD particles, more traditionally termed PSD.
When a SD particle becomes larger the magnetic moments tend to “flower out”in the corners of that par-
ticle. This state is termed flower state. In slightly larger particles the atomic spin magnetic moments curl
within the particle around one or more cores. This state is called vortex state. An ideal vortex has a low
magnetostatic energy because the particle's residual magnetic moment is low. Vortex cores may evolve
into classic magnetic domains in yet slightly larger grains.
•Magnetostatic interactions: When magnetic particles are located sufficiently close to each other, their mag-
netic fields will interact. The interaction field depends on the configuration of neighboring particles and
will be stronger if more adjacent particles are magnetized in the same direction. Magnetostatic interac-
tions are an important factor to explain the magnetic properties of closely spaced aggregates or intergrown
magnetic particles in earth materials.
•Magnetostriction: a property of ferromagnetic (sensu lato) materials which causes them to expand or con-
tract in response to an external magnetic field. It is considered to be a coupling effect between the mag-
netic energy and the mechanical energy observed in ferromagnetic materials under the influence of the
external magnetic field.
•Multidomain (MD): With increasing grain size, the magnetostatic energy of a magnetic particle increases.
To minimize this energy, a particle will begin to nucleate domain walls at a critical grain size threshold.
These walls divide the particle into two or more magnetic volumes or domains. The magnetization is uni-
form in each domain, but it differs in direction from domain to domain. The transition between small
grains with only one domain (termed SD) and MD grains is not sharp; a size range exists with a noncol-
linear spin structure that may contain a few domains.
•Natural remanent magnetization (NRM): The remanent magnetization that has been acquired naturally
(i.e., not artificially acquired in a laboratory). It is the remanent magnetization of a rock that is present
before any laboratory experiments are carried out. This preserves a record of the Earth's magnetic field
at the time the mineral was laid down as sediment or crystallized in magma. In many cases, that original
NRM component—termed primary NRM—can be retrieved from the composite NRM with paleomag-
netic techniques. Thus, with paleomagnetic data the original position of rock units can be reconstructed,
so their plate tectonic movement can be quantified.
•Néel temperature (T
N
): In antiferromagnetic substances, the individual magnetic moments are aligned
antiparallel to one another below a certain critical temperature, which is known as Néel temperature
(T
N
). In practice, Curie and Néel temperature are being used interchangeably. Named after the French
physicist Louis Néel (1904–2000).
•Paramagnetic: The magnetic state in materials with uncompensated electron spins that do not behave
collectively, that is each free electron spin is an individual atomic magnet. In an applied magnetic
field these moments will partially align to that applied field: A paramagnetic material will be magne-
tized in the direction of its inducing magnetic field. When the field is removed, the induced magnetic
moment disappears instantaneously. Paramagnetism is proportional to the inverse of the absolute
temperature.
•Pseudo single domain (PSD): A magnetic structure that is intermediate between the single‐domain (SD)
and MD states. A PSD particle contains more than one domain (up to ∼5–7) but exhibits many of the
stable magnetic properties typical of SD particles. See also magnetic vortex state.
•Saturation magnetization (M
s
), saturation remanent magnetization (M
rs
): see hysteresis loops.
•Single domain (SD): A uniformly magnetized magnetic particle with a single magnetic domain. In most
ferromagnetic and ferrimagnetic minerals, SSD grains are extremely small (the SD size range in equant
magnetite is ∼30–80 nm at room temperature).
•Superparamagnetism (SP): Class of magnetic behavior exhibited by very small particles (< ~25–30 nm in
magnetite at room temperature) with relaxation times (the magnetic decay time equivalent to the half‐life
of radioactive elements) smaller than laboratory timescales (typically taken at 100 s). For these particles,
atomic magnetic moments align in low, Earth‐like, applied magnetic fields to produce a strong induced
magnetization that is rapidly attenuated by thermal vibration after removal of the field (seconds to
minutes).
•Thermal remanent magnetization (TRM), also thermoremanent magnetization: The remanent magnetiza-
tion that is acquired by a sample when it is cooled in the ambient geomagnetic field, from above the Curie
temperature of the constituent magnetic minerals.
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A.2 Part II: Fault Rock and Faulting‐Related Terms
•Borehole breakout: Drilling‐induced compressive failures of the borehole wall occur once the stress
exceeds the strength of the rock or sediment. Their geometry is believed to replicate reliably the orienta-
tion and magnitude of in situ principal stress. Borehole breakouts are deemed to be key indicators of the
stress state in the subsurface. See Zoback et al. (2003) for more details.
•Brittle deformation: Deformation involving a “throughgoing discontinuity”in the rock. It includes tensile
cracking, shear fracturing, and frictional sliding. Brittle deformation occurs when the stresses on a rock
exceed the failure strength of the rock, causing a loss of cohesion.
•Cataclasite:Afine‐grained, cohesive fault rock that generally forms at shallow depths in the crust (typi-
cally from ~5 to ~10–15 km), dominantly by brittle deformation processes such as microcracking and
abrasion. It may be subdivided according to the relative proportion of finer‐grained matrix into protoca-
taclasite (with 10–50% matrix), mesocataclasite (with 50–90% matrix) and ultracataclasite (90–100%
matrix).
•Cataclastic flow :Distributed brittle deformation associated with motion of rock particles on a large collec-
tion of cracks and/or frictional surfaces.
•Clast: Fragments of rock surrounded by a finer‐grained matrix.
•Clastic dike: A seam of sedimentary material that fills in an open fracture and cuts across sedimentary
rock strata or layering in other rock types. Clastic dikes form rapidly by fluidized injection (mobilization
of pressurized pore fluids) or passively by water, wind, and gravity (sediment swept into open cracks).
•Damage zone: The region encompassing a mappable network of subsidiary brittle deformation structures
surrounding the fault core. These structures may include small faults, veins, fractures, cleavage, and folds
(Caine et al., 1996).
•Ductile deformation:“Continuous and homogeneously distributed”deformation. It can involve numerous
mechanisms ranging from macroscopic cataclastic flow to crystal‐plastic processes (Snoke et al., 1998).
Ductile deformation occurs when the stresses on a rock cause permanent strain, without the loss of
cohesion.
•Earthquake nucleation: The precursor stage preceding the dynamic rupture propagation of an earthquake
(with negligible inertial effects), which involves quasi‐static, aseismic slippage of a fault surface that has a
maximum scale termed “nucleation length.”
•Earthquake energy dissipation: During an earthquake, the elastic strain energy accumulated during the
interseismic period is released by rapid fault frictional sliding. The total energy released (E
tot
)is
partitioned into breakdown work (W
b
), frictional heat (E
H
), and radiated energy (E
R
), that is,
E
tot
=W
b
+E
H
+E
R
. Note that W
b
includes the fracture (surface) energy G, which is the energy required
to cause fracture near the tip of a fault during rupture, and the remainder is also transferred to frictional
heat. E
H
is the energy dissipated by frictional heating during slippage and is often considered to be the
largest component (~80–90%) of the total energy budget. Increasing temperature during coseismic slip
can trigger various chemical reactions and phase transitions. These processes are mostly endothermic,
but the taken‐up energy (often referred to as the EC) is usually minor (<1%) compared with frictional heat
generated.
•Fracture: A general term for any break in a rock mass, whether or not it causes displacement, including
cracks, joints and faults.
•Fault breccia: A coarse‐grained incohesive fault rock, formed by brittle deformation processes, consisting
of at least 30% visible fragments (>0.1 mm).
•Fault core: A zone (often asymmetric) enclosed by damage zones that results from the highly localized
strain and intense shearing that accommodates the majority of the displacement within the fault zone.
It generally consists of a number of recurring slip surfaces and fault rocks such as gouges, cataclasites,
and breccias.
•Fault gouge: An incohesive, clay‐rich fine‐to ultrafine‐grained cataclasite, which may possess a foliation
and contains <30% visible fragments. Lithic clasts may be present.
•Matrix: Groundmass of fine‐grained material, often surrounding larger fragments.
•Mylonite: A fault rock which is cohesive and characterized by a well‐developed foliation resulting from
tectonic grain‐size reduction derived from crystal plastic processes. It commonly contains rounded por-
phyroclasts and lithic fragments of similar composition to minerals in the matrix. Mylonites may be
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subdivided according to the relative proportion of fine‐grained matrix into protomylonite, mesomylonite
and ultramylonite, which feature grain‐size reduction in respectively <50%, 50–90%, and >90% of the rock
volume.
•Principal slip zone (PSZ): The fault plane along which the most slip is accommodated during a single
earthquake event. It is located within the fault core, surrounded by the damage zone. Synonyms: principal
slip surface (PSS) and principal slip plane (PSP).
•Pseudotachylyte: A cohesive fault rock that occurs as solidified melts produced by frictional heating during
seismic slip. It is typically dark in color and is glassy in appearance. Generally, quick, quench cooling of
the melt along the main fault plane forms a pseudotachylyte generation vein. Sometimes, some of the melt
may intrude minor faults or coseismic fractures into the wall rock; the resulting veins are termed injection
veins.
•Velocity‐strengthening/weakening: The fault strength in the brittle crust is a function of two basic para-
meters: the effective normal stress and the coefficient of friction. A fault may be strong and slip stably,
or weak and fail in large earthquakes. To explain differences between stable sliding (aseismic) and
unstable stick‐slip (seismogenic) behavior, frictional velocity dependence is considered the most likely
mechanism. Materials that exhibit velocity strengthening (frictional resistance increases with sliding velo-
city) produce only inherently stable frictional slip, whereas those that exhibit velocity‐weakening
(frictional resistance decreases with sliding velocity) frictional behavior are capable of hosting unstable
rupture.
Data Availability Statement
Data of Figure 4d can be found at Yang et al. (2018); data of Figures 6a–6d are available in Özdemir
et al. (1993, 2008), Dekkers et al. (1989), and Guyodo et al. (2006); and data in Figures 11b and 11c are from
Chou, Song, Aubourg, Song, et al. (2012) and Chou, Song, Aubourg, et al. (2014). The data used to create
Figure 5 can be found in the Magnetics Information Consortium (at https://doi.org/10.7288/V4/MAGIC/
16879).
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Acknowledgments
This work was supported by the
National Science Foundation of China
(NSFC) grant numbers 41874105,
41472177, and 41204062 to T. Y.
Y. M. C. was funded by the Shenzhen
Science and Technology Program under
grant KQTD20170810111725321,
Science and Technology Innovation
Committee of Shenzhen Municipality
under grant ZDSYS201802081843490,
and the Southern University of Science
and Technology under grants
K19313901 and Y01316111. This work
was also supported by the National
Science Council of Taiwan (grants NSC
101‐2116‐M‐003‐005 and NSC
102‐2116‐M‐003‐003 to E. C. Y. and
Y. M. C.) and National Science
Foundation (Instrumentation and
Facilities grant EAR‐0521558 and
Geophysics grant EAR‐0228818 to
E. C. F.). M. J. D. acknowledges support
from Netherlands Science Foundation
(NWO) Deep NL grant 2018.040. We
are grateful to Huan Wang and Tsafrir
Levi for providing the original copies of
Figures 2h and 19, respectively. We
thank Toshiaki Mishima for discussions
on an earlier version of the manuscript.
We also thank Silvia Mittempergher,
Bjarne S. G. Almqvist, and two anon-
ymous reviewers for their insightful
comments and suggestions that helped
to improve the manuscript signifi-
cantly. The Editor‐in‐Chief Fabio
Florindo is acknowledged for the effi-
cient handling of the manuscript.
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