Byrne et al.
Progress in Earth and Planetary Science(2024) 11:23
https://doi.org/10.1186/s40645-024-00627-w

Progress in Earth andPlanetary Science

Open Access

REVIEW

Tectonic exhumation of a metamorphic corein an arc‑continent collision during obliqueconvergence, Taiwan
Timothy Byrne1* , Michael Chojnacki1,2, Jonathan Lewis3, Jian‑Cheng Lee4, Gong‑Ruei Ho4,10, En‑Chao Yeh5,
Yuan‑Hsi Lee6, Chin‑Ho Tsai7, Mark Evans8 and Laura Webb9

Abstract
Observations over the last few decades from a number of orogenic systems have highlighted the possible impor‑tance of tectonic exhumation, i.e., ductile thinning and normal faulting, in exhuming rocks once buried in highpressure conditions. Taiwan is one of the few active orogens in the world where rocks that once experienced highpressure metamorphism (> 50 km) are exposed at the Earth’s surface, providing a natural laboratory for advancing ourunderstanding of exhumation processes. We integrate previously published studies of the Taiwan orogen with new
structural, geochronological, and fluid inclusion microthermometry data to argue that tectonic extrusion and struc‑tural thinning played a critical role in exhuming the metamorphic core of the orogen until very recently, ca. 0.7 Ma.
We propose a two-stage process for exhuming the high-pressure metamorphic rocks of the Yuli Belt: an initial stage
where exhumation is driven primarily by pressure gradients in a subduction channel and a second stage that is ini‑
tiated as an orogen-parallel regional-scale strike-slip zone, the Tailuko shear zone, is offset by an orogen-normal
strike-slip zone. The offset generates an extensional bend that is filled with extruding high-pressure rocks as the upper
crust is structurally thinned. Evidence for tectonic thinning comes primarily from a low-angle penetrative foliationthat records significant vertical shortening and a suite of sub-vertical late-stage, mineral-filled veins. Isotopic dating
indicates that the second stage started ca. 2.4–3.1 Ma and ended at ca. 0.7 Ma when the northern Backbone Range
orocline started to form. We propose the low-angle foliation formed in the footwall of a regional-scale extensional
shear zone that rooted to the east, beneath the forearc. Combined tectonic and erosional processes may have limited
the topographic growth of the orogen from ~3.0 to < 1.0 Ma.
Keywords Exhumation, Arc-continent collisions, Ductile thinning, Extrusion, Taiwan*Correspondence:
Timothy Byrne
Tim.byrne@uconn.edu
1
Department of Earth Sciences, University of Connecticut, 354 Mansfield
Road, Storrs, CT 06269, USA
2
USEngineering Solutions, Hartford, CT, USA
3
Department of Geography Geology Environment and Planning, Indiana
University of Pennsylvania, 114 Walsh Hall, Indiana, PA 15705, USA
4
Institute of Earth Sciences, Academia Sinica, Nankang, Taipei 115, Taiwan
5Department of Earth Sciences, National Taiwan Normal University,
Taipei 116, Taiwan
6Department of Earth and Environmental Sciences, National Chung
Cheng University, Chia‑Yi 621, Taiwan
7
Department of Natural Resources and the Environment, Nation Dong
Hwa University, Shoufeng 974, Taiwan

8Department of Earth and Space Sciences, Central Connecticut State
University, New Britain, CT 06053, USA
9
Department of Geography and Geosciences, University of Vermont,
Burlington, VT 05405, USA
10
Division of Geology, National Museum of Natural Science,
Taichung 404023, Taiwan© The Author(s) 2024. Open Access This article is licensed under a Creative Commons Attribution 4.0 International License, which
permits use, sharing, adaptation, distribution and reproduction in any medium or format, as long as you give appropriate credit to theoriginal author(s) and the source, provide a link to the Creative Commons licence, and indicate if changes were made. The images orother third party material in this article are included in the article’s Creative Commons licence, unless indicated otherwise in a credit line
to the material. If material is not included in the article’s Creative Commons licence and your intended use is not permitted by statutory
regulation or exceeds the permitted use, you will need to obtain permission directly from the copyright holder. To view a copy of thislicence, visit http://creativecommons.org/licenses/by/4.0/.

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

1 Introduction
Geologic processes that exhume metamorphic rocks to
the Earth’s surface include erosion, normal faulting, andductile deformation that thin the Earth’s crust. Erosion
is relatively well understood and documented in many
orogenic systems worldwide and, when integrated withappropriate material flow paths in the orogen, can be theprimary exhumation process (Malavieille 2010; Willett
et al. 2003). Observations over the last few decades froma number of orogens have highlighted, however, the possible importance of ductile thinning and normal faulting(Platt 1993; Grujic et al. 1996; Ring and Kumerics 2008;
Ring and Brandon 1999). These recent observations raiseseveral key questions about the relative importance oftectonic versus erosional processes in orogenesis.
The orogenic belt in Taiwan is well suited to better
understand tectonic exhumation because it is relatively
accessible, displays both late-stage normal faults and a
low-dipping ductile fabric, and is one of the few active
orogens in the world where rocks that once experienced
high-pressure metamorphism (> 50 km) are exposed at
the Earth’s surface (Jahn et al. 1981; Lan and Liou 1981;
Liou 1981). The nature of on-going and rapid exhumation of the metamorphic core of Taiwan, by either erosion or structural processes, is, therefore, fertile groundfor inquiry (e.g., Suppe 1981; Kirstein et al. 2009; Willett et al. 2003; Malavieille and Trullenque 2009; Dadson
et al. 2003).
Here, we integrate previously published studies of the
Taiwan orogen with new structural, paleotemperature,and geochronological data to argue that tectonic exhumation (e.g., extrusion and tectonic thinning) has been
under-appreciated in the Taiwan orogen and proposethat it played a critical role in exhuming the metamorphic core of the orogen until very recently, ~0.7 Ma.
Specifically, we focus on late-stage structures generally
associated with the ductile–brittle transition in the metamorphic core. The primary late-stage structures recognized in the core are a low-dipping penetrative foliationof variable intensity and mineral-filled veins (primarily quartz and calcite and locally adularia) that generally
occur perpendicular to the penetrative foliation. Ductile
normal-sense shear zones and normal faults lined with

Page 2 of 50

pseudotachylite are also included in the suite of late-stage
structures associated with the ductile to brittle transition.
The veins and low-dipping foliation have been recognized throughout the eastern Backbone Range, including the Tailuko and Yuli Belts and significant portionsof their sedimentary covers, the Backbone Slate Belt andthe Eastern Slate Belt, respectively. However, the ductile
normal faults and pseudotachylite appear restricted to
the northern Backbone Range. Structures that may haveformed in the shallow crust (< 3–5 km) above the ductile–brittle transition, for example, joints and zones of
localized brecciation or cataclasis, are not included.
We propose a two-stage process for exhuming thehigh-pressure metamorphic rocks of the Yuli Belt. During the initial stage, exhumation is driven primarily bypressure gradients in a subduction channel, possibly
enhanced by partial subduction of continental crust.
This process causes the rocks to be tectonically extruded
up the subduction channel with vertical and horizontal
components. The second stage is initiated as a marginparallel, left-lateral strike-slip system is offset, forming
a left-stepping extensional bend or pull-apart zone. The
extruding high-pressure rocks fill this extensional zone,
as the hanging wall is structurally thinned. The additionof tectonic processes in removing rock from the orogen
may explain the unusually fast rates of exhumation preserved in low-temperature thermochronological data,
starting at 2.5–2.0 Ma. Combined tectonic and erosional
processes may have also limited the topographic growthof the orogen from ~3.0 to < 1.0 Ma.

2 Tectonic setting
The development of the arc-continent collision in Taiwan is commonly described as a propagating collision
with the north-trending Luzon arc, sitting on the western edge of the PSP, as it moves toward ~315°, colliding
with the ENE-trending passive margin of Eurasia (Teng1990) (Fig. 1). These geometries and kinematics predict asouthwest propagating collision (Suppe 1984), suggestinga space–time equivalence along the strike of the orogen.
In this context, the orogenic system in the northeast isinterpreted to be older or more mature than the southwest system. Several authors have also used this unique(See figure on next page.)
Fig. 1 Tectonic and geologic settings of Taiwan. A Plate tectonic setting showing northwest convergence between the Philippine Sea and Eurasian
Plates (PSP and EUR, respectively) at ~7.8 cm/yr (yellow arrow). Gray triangle in northern Taiwan shows Ryukyu block (RYUK), which is rotating
CW ~47°/Myr, red arrow, based on GPS data (Rau et al. 2008). Yellow star represents subduction reversal inflection point. B Geologic mapof Taiwan (modified from Lin and Chen 2016). Units in color are discussed in the text. Note 50° change in orientation of cleavage and 48° change
in orientation of late-stage veins north of the orocline C relative to veins south of the orocline D is consistent with instantaneous CW rotationof the RYUK block and suggests rotation started ~1 Ma. OT: Okinawa Trough. RA: Ryukyu Arc. LV: Longitudinal Valley. Stereonets and all subsequent
nets are equal-area, lower hemisphere projections

A)(2024) 11:23

Page 3 of 50

B)

Taipei Basin

Mean (P/T)= 18/081

OT

30

Byrne et al. Progress in Earth and Planetary Science

30

60

0

Yilan PlainRYUK block;

W

E

RA

Myr

12

0

24

0

15

x.
pro
ap linecoroS

N = 35C) Poles to late-stage
veins north of orocline
Hep

PSP

0

21

0

EUR

ing

CXI

Kuk

H

utsu

Liw

R.

Cli

s

uR

Mug
Sho

ne R

ufen

.

gR

kbo

LV

Hon

Fig 5

CR

ulu

60

0

12

0

24

0

21

0

R.

E

0

gsu
e

Xinw

30

15

Chi

H

N

W

gye

SXI

.

Mean (P/T)= 22/033

30

ang

e

ua R

Bac

Foreland

.

S

i R.

D) Poles to late-stage
veins south of oroclineR.

Fig 13A
Shoufeng Fault

M
M= Vein orientation
data, blue and blacklled (this study)= Fluid inclusion
samples (this study= 40Ar/39Ar sample(this study= Mean cleavage (Yen,1970)
Fig. 1 (See legend on previous page.)

Zhib

M=Miocene, recently recognized

en R

Taim

Mt
Beidawu

ali R

.

.

N = 150

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

setting and the approximate balance between erosion andaccretion to argue that the orogen has achieved a steadysize and shape during the last few million years (Willett
et al. 2003; Suppe 1981; Fuller et al. 2006).
However, the propagating collision model for Taiwan,interpreted from map patterns and plate kinematics,
appears to mask a more complicated collision involving
flipping subduction, back-arc spreading, and orogenic
oroclines. For example, in the south, the Eurasian Plate
dips east and subducts beneath the Philippine Sea Plate,whereas, in the north, the Philippine Sea Plate dips northand subducts beneath the Eurasian Plate and the northern Backbone Range of Taiwan. Refraction data (Van
Avendonk et al. 2015; Ustaszewski et al. 2012) and theidentification of subducted slabs using global tomography (Wu et al. 2016) show the Eurasian lithosphere dipping approximately 60° beneath southern Taiwan andnearly 90° beneath northern Taiwan. Patterns of seismicity and earthquake focal mechanisms in northern Taiwan
also suggest that the PSP is impinging on the subducted
Eurasian Plate (Lallemand et al. 1999, 2013; Ustaszewski
et al. 2012). Together, these data argue for a reversal in
the subduction polarity with the inflection point in central Taiwan (von Hagke et al. 2016) (Fig. 1).
The orientation of the plate boundary relative to theconvergence vector also changes dramatically from south
to north, adding to the complexity of the orogen. In thesouth, the plate boundary is at a high angle to the vector, which trends ~315°. In the north, however, the plateboundary trends east–west along the Ryukyu Arc and is
at a much lower angle to the convergent vector. Subduction of the Philippine Sea Plate beneath the Ryukyu Arc
has also generated back-arc spreading in the Okinawa
Trough, starting as recently as 1.5 Ma (Fig. 1) (Osozawa
et al. 2012; Sibuet et al. 1998, 2021). After a period of relative quiescence, ending ~1.0 Ma, extension in the backarc propagated westward to Taiwan (Sibuet et al. 1995).
This active rift forms the Yilan Plain in northern Taiwan,
separating the northeasterly trending Hsüehshan Range
from a more easterly trending Backbone Range (e.g., seechange in strike of foliations in Fig. 1).
GPS data north and south of the Yilan Plain are also
consistent with propagation of back-arc spreading
between the two ranges and clockwise rotation of thenorthern Backbone Range, forming an active orogenic
orocline (Fig. 1B) (Rau et al. 2008; Hu et al. 2002). Rau
et al. (2008) used GPS data and a block rotation model(Mccaffrey 2002) to define 11 blocks in northern Taiwan,
with the “RYUK” block representing the eastern limb of
the orocline (Fig. 1). The block model shows the RYUK
block rotating clockwise 47° +/− 7.8°/Myr. Using thismodern rotation rate and the ~50° difference in strikesof the southern and northern limbs of the orocline, i.e.,

Page 4 of 50

the central and northern extents of the Backbone Range(Fig. 1), we propose that the orocline started to form
~1.0 Ma. This age of orocline initiation is consistent withthe transition from compression to extension in northern
Taiwan at ~0.8 Ma (Teng et al. 2001; Lee and Wang 1988)
and the initiation of sediment deposition in the Taipei
Basin at 0.4 Ma, an analogous basin to the Yilan Plain(Wei et al. 1998; Teng et al. 2001).
The development of the orocline in the northern Backbone Range < 1 Ma also provides critical temporal constraints on the structural fabrics that occur within andoutside of the orocline. The orocline is particularly conspicuous on regional-scale geologic maps of Taiwan,
where the strike of the dominant fabric in the Tailuko
Belt changes from ~020° to ~070° in the northern partof the range (Fig. 1). A low-velocity crustal root in thenorthern part of the range mimics the surface geology(Huang et al. 2014), suggesting that the orocline is a crustal-scale feature. In more detail, Ho et al. (2022) proposedthat the dominant fabrics in the Tailuko Belt, defined bya penetrative foliation, ­S2, and lineation, ­L2, follow themap-scale pattern and, therefore, appear to be deformedby the orocline. This interpretation is consistent with40
Ar/39Ar ages of syn-kinematic biotite grains from a S­2mylonite zone that range from 4.1 to 3.0 Ma (Wang et al.
1998). In the following sections, we follow a similar logic
and propose that late-stage structures in the metamorphic core formed between 3.0 and 1.0 Ma because they
cross-cut ­D2 in the Tailuko Belt and are deformed by the
orocline. For example, the orientation of late-stage veins
north of the orocline axis shows an apparent clockwise
rotation of 47° relative to late-stage veins south of the axis
(Fig. 1).

3 Geologic backgroundThe steady subduction of the Eurasian lithospherebeneath the Luzon Arc resulted in the progressive imbrication and accretion of rocks with oceanic and continental affinities, forming the Taiwan orogen (Fig. 1) (Ho
1988; Teng 1990). The Tailuko Belt, characterized primarily by phyllites, schists, marble, and granitic bodies,and the Yuli Belt, characterized partly by high-pressure
metamorphism, represent the metamorphic core of theorogen. These belts, however, have strikingly different
depositional, structural, and metamorphic histories andare separated by the enigmatic Shoufeng Fault (Fig. 1)(Yen 1963; Ho and Lo 2015). The Tailuko Belt is primarily
Mesozoic and is overlain by an Eocene to Late Miocene
sedimentary cover. The dominant penetrative foliation
strikes parallel to the orogen and appears to display acleavage fan with an east-dipping slaty cleavage in thewest and a west-dipping crenulation cleavage in the east.
In contrast, the Yuli Belt is Miocene, and its subunits,

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

locally mapped as members, form gently north-dipping
tectonostratigraphic units with complex internal structures. The units may represent stacked thrust sheets, as
they appear to pinch out in cross section or are separated
by discontinuous lenses of mafic to ultramafic rocks (Lin
1999b; Tan et al. 1978). Both belts of rocks, however, display a suite of late-stage structures that appear to “stitch”
the units at the time of their development.
3.1 Tailuko belt and its sedimentary cover

Previous structural studies of the Tailuko Belt and its
cover, the Slate Belt of the central Backbone Range,
suggested three episodes of deformation during NWdirected convergence, generalized here as D1, D2, and D3(Lin et al. 1984; Faure et al. 1991; Clark et al. 1992; Hoand Lo 2015; Stanley et al. 1981; Lu and Wang Lee 1986),
although Stanley et al. (1981) recognized three additional
deformation events. The early deformation episodes, D1
and D2 (and D3 and D4 of Stanley et al. 1981), included
the development of an associated penetrative foliation
and a variety of outcrop-scale folds interpreted to berelated to the accretion of a passive margin sequence(Lin et al. 1984; Faure et al. 1991; Stanley et al. 1981). The
general west dip of the dominant foliation in the easternpart of the belt also motivated interpretations of eastdirected thrusting (see, e.g., Ernst 1977; Page and Suppe
1981; Faure et al. 1991). Stanley et al. (1981) elaborated
on this concept and proposed that his D3 and D4 events
formed during the development of a regional-scale, eastverging fold. Most subsequent workers followed Stanley
et al. in explaining the west-dipping cleavage in the eastern part of the belt (e.g., Lin et al. 1984; Lu and Wang
Lee 1986; Faure et al. 1991). In addition to the change in
dip, the regional-scale fold interpretation was based onlithostratigraphic correlations that assumed a Mesozoic
age for the Yuli Belt and, therefore, its approximate contemporaneity with the Tailuko Belt. Considering that theYuli Belt is Miocene rather than Mesozoic (Mesalles et al.
2020), the lithostratigraphic correlations and regionalscale folding merit reevaluation.
More recently, Ho et al. (2022) redefined D2 in the
Tailuko Belt and documented evidence of significant
regional-scale, left-lateral ductile shear. In addition, Ho
et al. (2022) proposed that strike-slip deformation in the
Tailuko Belt was contemporaneous with deformation in
the Slate Belt, which crops out west of the Tailuko Beltand records west-directed thrusting. They propose thatthe two belts formed as highly oblique plate convergence
was partitioned into plate-boundary parallel strike-slipin the Tailuko Belt and plate-boundary normal thrusting and imbrication in the Slate Belt to the west closer
to the trench. Integrated geochronologic data suggeststhat partitioning was active from ca 6 to 3 Ma (Ho et al.

Page 5 of 502022). The inactivity of ­D2 in the Tailuko Belt at ~3 Ma isconsistent with the initiation of the orocline in the northern Backbone Range at ~1 Ma because D2 mylonites arerotated clockwise by the orocline.
A suite of structures that cross-cut S2 and appear torepresent a transition from ductile to brittle deformation have been recognized in the Tailuko Belt (Fig. 2)(Table 1), and a few of the studies suggest a transition
from horizontal to vertical shortening. For example, outcrop-scale folds with gently dipping axial surfaces, oftenwith a semi-penetrative, axial planar pressure solution
cleavage, have been mapped along the South and Central Cross-Island Highways (SXIH and CXIH, respectively) (Clark et al. 1992; Faure et al. 1991; Stanley et al.
1981; Lu and Wang Lee 1986; Lee 1995) (Fig. 2). Lee andYang (1994) also documented a set of late-stage quartz
veins along the CXIH that record northeast-southwest
extension. More recently, Ho et al. (2022) and Ho andLo (2015) completed a detailed map of the Tailuko Belt
along the Shoufeng River and showed that a gently dipping foliation cross-cuts S2, the dominant foliation in thearea, resulting in the formation of outcrop-scale recumbent folds of S2. They also proposed that the folds andgently dipping foliation formed contemporaneously withhorizontal extension. Ho and Lo (2015) labeled the foliation “S3” (Fig. 2, Site E), consistent with previous studiesof the Tailuko Belt (Clark et al. 1992; Faure et al. 1991; Lu
and Wang Lee 1986; Lee 1995), excluding Stanley et al.
(1981). Stanley et al. recognized three additional deformation events, and, based on their descriptions and ourfieldwork along the South Cross-Island Highway, we correlate D3, as documented by Ho and Lo (2015), to D6 in
Stanley et al. (1981).
Crespi et al. (1996) mapped suites of late-stage normal
faults along the South and Central Cross-Island Highways, and showed that the faults locally involved crystal plastic and brittle deformation. Kinematic analyses
yielded east–west extension along the CXIH and NE-SW
extension along the SXIH (Fig. 2 Site D and D’, respectively). They also proposed that late-stage extension in
the eastern Tailuko Belt along ease-dipping shear zones
caused clockwise rotation of the older fabrics, providing
an alternative interpretation for the cleavage fan in the
Tailuko Belt.
In the northern Tailuko Belt, pseudotachylite-lined
normal faults and adularia-filled veins cross-cut thedominant foliation, S2, and record east–west extension (Korren et al. 2017; Chen et al. 2022), providing
temporal and kinematic constraints on the late-stage
deformation (Fig. 2, Site B). 40Ar/39Ar step-heating
analyses of the pseudotachylite and adularia also yieldnearly identical dates; 1.55 +/− 0.5 Ma for the pseudotachylite (Chen et al. 2017a) and 1.09 +/− 0.03 Ma andByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 6 of 50

A)

N
30

0

N

0

60

0
24

12

W

E

0

0

21

15

0

B)

x.
pro
ap line
c
oro

30

33

C)

S

D)

CXI

60

0

W

0

0

33

W

gR

12

0

21

33
0

30

0

21

15

0

0

S

30

60

0

2.6 MaJ)

ChaochouFault

3.6 MaK) Zhibe

nR

.

R.

E
12

0

0

24

0

ulu

3.4 Ma

i R.

15

Xinw

W

sue

0

21

0

hing

SXI
H

21

I) C

15

0

0

S

Sc

N

R.

0

gye

33

Hon

12

0

E

.

G)

E

24

60

0

24

U
W

.

ufen

(F)

D

60

0

N

30

0
15

Sho

Tayuling
Fault (ref. M)

N

.

ua R

0

S

30

uR

Mug

12

0

24

H)

Liw

H

E

30

0

30

33

30

30

E)

N

S

Lc

S1

2.4 Ma3.1 Ma= AFT ages (ref. L)

N)

L1

Fig. 2 Summary of previously published structural data interpreted to represent the latest penetrative event in each area. Two northern sites areon the RYUK block, which is rotating clockwise (Rau et al. 2008). Double arrows show extension or stretching directions; single arrows show motionof hanging wall. Geologic map from Chen, 2016. See Table 1 for explanations and references

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 7 of 50

Table 1 References and descriptions for Fig. 2
References

Description(A)

Yeh et al. (2016)

Black great circle and arrow show mean orientation of ductile shear zones and mean slip direction of hang‑
ing wall, respectively. N = 15. Site is on the RYUK block and is inferred to have rotated ~50° CW(B)Korren et al. (2017)Black great circle and arrow show mean orientation faults lined with pseudotachylite and mean slip directionof hanging wall, respectively N = 21. Site is on the RYUK block and is inferred to have rotated ~50° CW(C)

Lee (1995), Lee and Yang (1994) Black great circle: mean orientation of late fold axial surfaces from 10 sites along Central Cross-Island Highway(CXIH) and the Tailuko coastline. Black dots: poles to F3 fold axial surfaces. N = 66. Black arrows: stretchingdirection from quartz veins. N = 100(D) (D’) Crespi et al. (1996)

Black arrows: extension direction interpreted from late-stage normal faults along the CXIH and SXIH

(E)

Ho and Lo (2015)

Black great circle: mean orientation of S3 from 3 domains along the Shoufeng River. N = 83

(F)Yi et al. (2015)

Black arrow shows antiformal structure formed by S3 in the Yuli Belt(G)

Lin (1999a, b), Yen (1970), Lin
and Chen (2016)

Short dashed lines: intraformational strike-slip faults. Left-lateral faults from Lin (1999a; b) and Lin and Chen
(2016). Right-lateral strike-slip fault from Yen (1970)(H)

Clark et al. (1992)

Black great circles: mean orientation of S3 from 2 domains along the South Cross-Island Highway (SXIH). N =
276

(I)

Mondro et al. (2017)

Poles to composite foliation, Sc, and stretching lineation Lc, from Chulai and Hsinkao Formations (area in box).
N = 46 and N = 40, respectively(J)

Huang et al. (2014)

Regional-scale, late-stage antiform that deforms tectonic cleavage(K)

Conand et al. (2020)

Poles to foliation, S1, and stretching lineation, L1, from sedimentary rocks south of southern most exposuresof the Yuli Belt (area in box). N = 240 and N = 59, respectively(L)

Fuller et al. (2006)

Apatite fission track ages of totally reset detrital grains; pooled ages

(M)

Lee et al. (2022a, b)

Major out-of-sequence thrust that carries the metamorphic core(N)

Usami and Matsumoto (1940)

Late-stage left-lateral strike-slip faults on east and west sides of Tailuko Belt

1.5 +/− 0.3 Ma for two samples of the adularia (Wang1998; Chen et al. 2022). Korren et al. (2017) estimatedthe formation of the pseudotachylite at a depth of> 4 km with possible temperatures ranging from 250°
to 300 °C.
Further north in the sedimentary cover of the Tailuko
Belt, Yeh et al. (2016) mapped a suite of late-stage, ductile shear zones with normal displacements (Fig. 2, Site
A). The shear zones cross-cut ductile shear fabrics thatrecord a left-lateral sense of shear that we suggest correlate with ­D2 structures documented throughout the
Tailuko Belt (Ho et al. 2022). The younger, normal shear
zones also strike north–south, indicating east–west
extension, consistent with the pseudotachylite-lined
faults and adularia-filled veins in this area. We, therefore,
propose the normal-sense shear zones (Yeh et al. 2016)
formed contemporaneously with the dated extensional
structures.
The suite of structures described from the northern
Tailuko Belt that record east–west extension crop out inthe northern Backbone Range where regional structural
fabrics strike east-northeast and represent the northernlimb of the Backbone Range orocline (Fig. 1). The rocks
are also part of the RYUK tectonic block, which is rotating clockwise ~47°/Myr (Fig. 1, inset) (Rau et al. 2008).
Considering the rate of rotation, the curvature of theorocline, and the recent onset of extension in northern
Taiwan, we reconstruct these structures to their originalorientations with a 47° counterclockwise vertical-axis
rotation.
3.2 Yuli belt

The Yuli Belt is characterized by the abundance of metamorphosed clastic sedimentary sequences and the general absence of thick to massive marble units comparedto the Tailuko Belt (Wang Lee 1979; Lin et al. 1984; Yen
1963). Tectonic or sedimentary blocks of pillowed lava,
mafic and ultramafic bodies, and associated serpentinites
and plagiogranites also characterize the Yuli Belt (Yuiet al. 1990; Wang Lee and Lu 1984; Lin et al. 1984; Lan
and Liou 1981, 1984); many of these blocks also record
metamorphic conditions suggesting up to 50 km of burial(Yui et al. 1996, 2009, 2012; Huang et al. 2023; Beyssac
et al. 2007; Tsai et al. 2013; Keyser et al. 2016; Baziotis
et al. 2016). Early work suggested the belt was Mesozoic
and possibly contemporaneous with the Tailuko Belt
(Yen 1963; Stanley et al. 1981; Ho 1986, 1988). However, U–Pb ages of detrital zircons from the belt indicate
a maximum depositional age of < 15.6 ± 0.3 Ma (Chenet al. 2017b). This new age is consistent with metamorphic ages from the Yuli Belt that range from 12.2 ± 0.5 Ma(Lo and Yui 1996; Chen et al. 2019) for 40Ar/39Ar ages
on amphiboles to 5.1 ± 1.7 Ma for Lu–Hf ages on garnet (Sandmann et al. 2015). The high-P rocks were thenexhumed to the present position starting from ~5–6 Ma(Chen et al. 2019; Huang et al. 2006; Lee et al. 2022b) andByrne et al. Progress in Earth and Planetary Science(2024) 11:23

experienced retrograde evolution as recorded by a phlogopite 40Ar/39Ar age 4.4 ± 0.1 Ma (Lo and Yui 1996) anda zircon U/Pb age of 3.3 ± 1.7 Ma from a clinozoisite veinin nephrite (Yui et al. 2014). Consequently, the Yuli Belt
represents one of the world’s youngest and most rapidly
exhumed high-pressure belts, with exhumation rates up
to 8 km/Myr.
Several field-based studies have outlined the general
structural patterns in the Yuli Belt, although detailed kinematic analyses are more limited. For example, Lin et al.
(1984) and Lin (1999b) proposed a tectonostratigraphy
composed of four gently north-dipping units (“Members” in Lin 1999b) based on field and petrologic datafrom transects along the Chingsuei and Wanjung Rivers and the area north and south of the Chouchi Fault(Fig. 2, site G). Although all of the units record protracted
deformation and metamorphic histories and rarely preserve stratigraphic markers, one of the structurally lowest
units, a “spotted” pelitic schist interlayered with greenschists and Mn-rich garnet schists, forms a distinctive
regional marker unit throughout the extent of the belt
(Yen 1963; Lin et al. 1984; Lin 1999b). Blocks of metamorphosed serpentinite, epidote amphibolite, metaplagiogranite, and metamorphosed pillowed volcanic rocks
are also often associated with this unit, suggesting an
oceanic origin for the belt (Lin et al. 1984; Liou 1981;
Lan and Liou 1984; Lo et al. 2020). Two higher structuralunits and one lower unit are richer in quartz, including
quartz and mica-quartz schists and quartzite, and areconspicuous in the absence of exotic blocks. The proposed tectonostratigraphy is similar to the sequences
observed in the Guangfu quadrangle (Yi et al. 2012) inthe central part of the Yuli Belt (Fig. 2, site F).
Lin et al. (1984) also completed one of the first detailed
structural studies of the Yuli Belt and recognized three
deformation stages: D1, D2, and D3. Each stage is associated with a suite of folds and foliations, with subsequentfolding events producing distinctive interference patterns. However, none of the deformation stages in theYuli Belt record significant evidence of left-lateral strikeslip faulting that might correlate with ­D2 structures
observed in the Tailuko Belt (Ho et al. 2022). Instead, thetectonostratigraphic units in the belt typically form gently, northwest to northeast dipping, structurally boundsheets, suggesting tectonic stacking. The distinctive“spotted schist” unit also appears to pinch out along andacross strike (Lin 1999b), consistent with tectonic stacking. Field relations, although limited, also show the exotic
blocks as lens-shaped clusters that parallel unit boundaries (Lin 1999b; Lin et al. 1984; Yen 1983; Tan et al. 1978).
In some locations, however, the exotic blocks appear
locally conformable with surrounding units (Lin et al.
1984) or are mapped as klippe sitting on units mapped

Page 8 of 50

as Yuli and the lower-grade sediments (Lin 1999b; Yen
1983).
S3 appears to have formed relatively late in the deformation history of the Yuli Belt (Lin et al. 1984). Alongthe Chingsuei River, ­S3 deforms S2, creating a crenulation cleavage or outcrop-scale recumbent folds with gently dipping axial surfaces. S3 also post-dates the growthof albite porphyroblasts and is associated with elongate
quartz, suggesting more limited ductile deformation thanassociated with the development of S2 (Lin et al. 1984).
Mapping in the central part of the Yuli Belt shows thefour tectonostratigraphic layers of the Yuli Belt forming
the western half of a north plunging antiform (Yi et al.
2012), and cross sections show the dominant structural
fabric, S2, overprinted by a gently dipping cleavage interpreted as S3 (Yi et al. 2012, 2015). S3 also forms a broad
antiformal structure that mimics the antiformal structure
defined by the tectonostratigraphy, although the interlimb angle is larger, ranging from ~146° to 170° (Fig. 2,
site F).
S3 is also locally deformed by a mylonitic fabric,
exposed along the western edge of the Yuli Belt in central Taiwan (Fig. 2, site F). Yi et al. (2015) completed adetailed geochronologic study across the mylonite zone
using 40Ar/39Ar on fine-grained muscovite grains thatform part of the mylonitic foliation and obtained six age
populations ranging from 0.68 to 1.16 Ma. The youngest
ages come from the middle of the mylonite zone, whereas
an older age of 2.65 Ma was obtained from the Yuli Belt
about 5 km east of the mylonite zone. The mylonite zone,
therefore, appears to have been recently active, although
the kinematics of this recent motion remain undefined.
However, a late-stage normal fault (Wangwutashan Fault,
Yi et al. 2015) crops out a few hundred meters west of
the mylonite, suggesting the possibility that the mylonite
formed during normal displacements. In any case, themylonite deforms and locally transposes S3 in the Yuli (Yiet al. 2015), constraining S3 to be older than ~0.7 Ma.
Transverse strike-slip faults also developed late in the
structural history of the Yuli and Tailuko Belts and may
have played a role in their exhumation. For example,
detailed mapping north and south of the northweststriking Chouchi Fault (Fig. 2, site G) indicates obliqueleft-lateral, normal slip across the fault (Lin 1999a). Lin(1999a) interpreted the net slip to be a few km, with thenorthern block moving up and to the west relative to thesouthern block. This interpretation is consistent withobservations across a geometrically similar strike-slip
fault north of the Chouchi Fault (Lin and Chen 2016),
suggesting that the two faults form a right-stepping, leftlateral fault system (Fig. 2) that crosses the metamorphic
core of the orogen. A third transverse structure westof the Chouchi Fault strikes approximately east–west

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23(Fig. 2), and Yen (1970) argued for right-lateral displacement based on contrasting structural geometries oneither side of the fault. The geometries and kinematics ofthe three faults are consistent with formation as a conjugate set during west-northwest shortening, althoughadditional research is needed to show their contemporaneity. In any case, they document late-stage shortening
and northeast elongation of the metamorphic core.
3.3 Low‑grade sedimentary rocks adjacent to the Yuli belt

The low-grade metasedimentary rocks exposed eastand south of the Yuli Belt also preserve late-stage structures that may correlate with structures observed in
the higher-grade metamorphic rocks. These rocks
were initially interpreted as Eocene (Fig. 1) (Stanley
et al. 1981). Recent U–Pb analyses on samples from thenorthern extent of the metasedimentary rocks, mapped
as the Chulai Formation (adopted from Stanley et al.
1981), however, yielded a maximum depositional age of~11 Ma (Mesalles et al. 2020), consistent with interpretations that the Chulai was in depositional contact with
the Yuli (Stanley et al. 1981). Stanley et al. (1981) also
proposed that the Chulai Formation was overlain by the
Hsinkao Formation south of the town of Yuli. We infer,
therefore, that the Hsinkao Formation, which forms thesouthern extent of the metasedimentary rocks, is also
Miocene.
The unconformity between the Chulai and Yuli Beltsuggests the Chulai may represent a relatively low-grade
sedimentary cover to the Yuli (Stanley et al. 1981), andrecent mapping of the metasedimentary rocks eastof the Yuli Belt suggests they form part of the Yuli Belt
(Chen et al. 2017b). Detailed mapping of the unconformity along the north bank of the Xinwulu River (Fig. 1)(Stanley et al. 1981) also revealed an outcrop-scale, eastverging asymmetric fold (see Fig. 5 in Stanley et al. 1981).
Although Stanley et al. (1981) related the fold to regionalscale back (i.e., east-directed) folding, the asymmetry andassociated structures are also consistent with the topto-the-east shear proposed by Crespi et al. (1996). An
alternative interpretation is that the sediments east andsouth of the Yuli Belt represent higher structural levels of
a middle to late Miocene accretionary prism (Chen et al.
2017b). We consider this possibility in more detail in the
Discussion.
A quantitative strain study of 18 samples from themetasedimentary rocks east of the Yuli Belt, including the Chulai and Hsinkao Formations, shows two
stretching events (Fig. 2, site I) (Mondro et al. 2017).
Incremental strain histories from syntectonic fibers instrain fringes around pyrite framboids were measured
relative to the orientation of the dominant cleavage,
interpreted to be associated with multiple strain states,

Page 9 of 50

and, therefore, described as a “composite” cleavage.
The composite cleavage in the belt dips gently northeast with a pi-diagram showing a broad, gently northeast plunging antiform (Fig. 2, site I). The composite
lineation also plunges gently northeast parallel to theregional fold axis. Because the strain history was noncoaxial, incremental strains were measured relative tothe dip of cleavage in cleavage-perpendicular sections
and relative to the strike of cleavage in cleavage-parallel sections. The results show two stages of progressive
extension: an early stage that is generally down dip anda later stage that is generally along strike. In geographic
coordinates, the early stage of extension trends generally northwest-southeast, whereas the later extension
trends northeast-southwest. Progressive finite stretch
values for the late-stage extension (i.e., data collectedfrom cleavage-parallel sections) typically range from 3
to 5, with a few as high as 14 (Mondro et al. 2017). The
earlier extension stage records a slightly higher maximum stretch, ranging from 3 to 11, with a few samples
below 3 and one sample recording a finite strain magnitude of 33 (Mondro et al. 2017).
One of the first detailed structural studies of the lowergrade sedimentary rocks south of the Yuli Belt suggeststhat they record three stages of deformation: D1, D2, andD3 (Lu and Lee 2001). The two early stages of deformation are interpreted to be related to west-directed shearing followed by east-directed shearing (Lu and Lee 2001;
see also Chang et al. 2009), whereas D3 includes a gently
dipping penetrative cleavage that overprints D1 and D2
structures. Strain fringes associated with D1 and D2 showa change in stretching directions from generally ENEWSW to NNW-SSE through time (Lu and Lee 2001). A
similar rotation in stretching directions, based on veinsets rather than strain fringes, is also recorded in a more
limited area along the coast (Shan et al. 2014), consistent
with the identification of late-stage north–south stretching in this area (Chang et al. 2009).
Finally and most recently, extensive mapping of deformation styles along the southern extent of the eastern
Backbone Range, including the areas mapped by Changet al. (2009), Shan et al. (2014), and Liu et al. (2001), suggests three structural domains with a central domain that
includes the metasedimentary rocks south of the Tailuko
and Yuli Belts (Fig. 2 site K) (Conand et al. 2020). The
metasedimentary rocks in the central domain (Domain II
of Conand et al. 2020) are characterized by gently southwest-dipping foliation and an associated down-dip (i.e.,
SW trending) stretching lineation (Fig. 2 site K). Extensional shear bands, S–C structures, and asymmetricpressure shadows in the central domain show increased
non-coaxial strain and consistently indicate a top-to-theSW sense of shear (Conand et al. 2020; Lu and Lee 2001).

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Liu et al. (2001) also suggested that top-to-the-SW shear
was associated with regional-scale left-lateral shearing.
Overall, the relation between the structures recognizedby different studies in the low-grade sedimentary rocks
appears complex. The rocks, however, appear to sharea change from generally EW to NS stretching, a top-tothe-SW sense of shear associated with a gently southwest-dipping foliation, and a possible role for left-lateral
strike-slip deformation.
3.4 Shoufeng fault

The Shoufeng fault generally represents the contact
between the Tailuko and Yuli Belts (Fig. 1) but documenting and understanding the origin and slip history of
the fault has been challenging. The fault is consistently
recognized as a zone of concentrated strain (Stanley et al.
1981; Yen 1963; Yi et al. 2012; Yui et al. 2012; Ho et al.
2022; Chen et al. 2017b), and early work suggested thatit originated as a low-angle, west-verging thrust (Stanley
et al. 1981). However, the present west dip of foliationsin the high-strain zone also motivated interpretations
involving east-verging thrust faults (Ernst 1977; Page andSuppe 1981; Liou et al. 1977) or folds (Stanley et al. 1981;
Faure et al. 1991; Lin et al. 1984). The west dip may also
reflect clockwise rotations (viewing north) of an initiallyvertical shear zone, as proposed by Crespi et al. (1996)
and Pulver et al. (2002).
Along the SXIH, detailed mapping of the boundary
between the Tailuko and Yuli Belts suggests at least twostages of deformation (Stanley et al. 1981). The dominant foliation, S2, dips moderately west and appears tocorrelate with the main foliation recognized throughout
the Tailuko Belt (Ho et al. 2022). This foliation is typically mylonitic and associated with left-lateral strike-slip
deformation (Ho et al. 2022). S2 is cross-cut by a moregently west-dipping foliation associated with west-verging asymmetric folds. The asymmetry indicated by thefolds is consistent with the proposal that the Shoufeng
fault is an overturned thrust zone that, before overturning, carried the Yuli Belt westward above the Tailuko
Belt (Stanley et al. 1981). The Shoufeng fault, therefore,
appears to record both strike-slip and thrust deformation, at least in the exposures along the SXIH.
In central Taiwan, the Daguan shear zone dips moderately west-northwest, separating the Tailuko and Yuli
Belts (“Takuan shear zone” in Yi et al. 2012)(Fig. 2 siteF), and may preserve deformation associated with theShoufeng Fault. However, the relatively young muscovite
ages obtained from mylonites in the shear zone, ranging
from 0.68 to 1.16 Ma, suggest, it records only the youngest stages of deformation associated with the fault. The
kinematics of this recent motion and the Daguan shear
zone are undefined, although the nearby Wangwutashan

Page 10 of 50

Fault is parallel to the Daguan shear zone and records
normal displacement (Yi et al. 2012).
Finally, recent mapping and map compilations suggestthat projections of the Shoufeng Fault and the left-lateral
strike-slip shear zones in the Tailuko Belt are poorly constrained south of the latitude of the SXIH (Fig. 1). For
example, most geologic maps show the Tailuko and Yuli
Belts and the Shoufeng Fault that separates the two belts,
continuing to the south-southwest, at least as far south
as the Zhiben River (Fig. 1) (Ho 1988; Chen 2000; Lin
and Chen 2016). However, more recent mapping shows
the fault curving southeast, cutting across the low-grade
sedimentary rocks to the Longitudinal Valley where thefault is apparently covered by the Quaternary sediments(Chen et al. 2017b). This interpretation suggests that
the Chulai and Hsinkao Formations (Stanley et al. 1981)
are part of the Yuli Belt. Similarly, recent map compilations (Lin and Chen 2016; Chen 2000) do not show two
regional-scale, left-lateral strike-slip faults mapped along
the east and west sides of the southern extent of the belt
(Fig. 2 site N) (Usami and Matsumoto 1940). The faults
appear to record several kilometers of offset (Usami andMatsumoto 1940), consistent with observations from the
Tailuko Belt that show significant left-lateral shear (Hoet al. 2022). These map-scale inconsistencies related to
the Shoufeng Fault and the Tailuko and Yuli Belts provide
rich opportunities for future research.
3.5 Exhumation cooling of the metamorphic core

Several studies of exhumation cooling in the Tailuko Beltand its sedimentary cover suggest progressively higher
exhumation rates since peak metamorphism about 6 Ma
(Beyssac et al. 2007; Lee et al. 2022b; Liu et al. 2001). Oneof the first age-elevation profiles from the belt suggested
two periods of accelerated exhumation cooling, from ~3
to 2 Ma and < 1 Ma (Liu 1982). The early cooling period
is consistent with thermally reset microcline grains thatyield 40Ar/39Ar plateau ages of 1.7 and 1.6 Ma (Lo andOnstott 1995). Although these studies were limited to thenorthern Backbone Range, four additional age-elevation
profiles from along the length of the range show a broadly
similar pattern with an initial acceleration in cooling at~2 Ma and a second acceleration < 1 (Hsu et al. 2016; seealso recent compilation by Lee et al. 2022b).
The early acceleration in exhumation was also confirmed through an integrated petrographic study ofadjacent sedimentary basins that show sediment accumulation rates increasing at ~2 Ma (Chen et al. 2019).
The second acceleration of cooling < 1 Ma (Liu 1982; Hsu
et al. 2016) has also been substantiated in the northern
Tailuko Belt through an additional age-elevation profile
incorporating subsurface data (Shen et al. 2020).

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Exhumation cooling in the Yuli Belt also appears tohave been relatively rapid (Chen et al. 2017b). Although
low topographic relief limits constructing age-elevation
profiles, totally reset fission track and (U-Th)/He ages of
detrital zircon and apatite grains from the eastern Backbone Range, including the Yuli Belt, the eastern part of
the Tailuko Belt, and the metasedimentary rocks southof the Yuli Belt, are < 2.0 Myr, indicating rapid coolingvery recently (see reviews by Beyssac et al. 2007; Lee et al.
2022b; Simoes et al. 2012; Fuller et al. 2006; Chen et al.
2019). These relatively young ages from multiple lowtemperature geochronologic systems occur in both the
Tailuko and Yuli Belts, suggesting that these belts preserve similar cooling histories since at least ~2.0 Ma (seealso Chen et al. 2019).
Low-temperature geochronologic data from the southern Backbone Range south of the exposures of the Tailuko and Yuli Belts provide additional constraints on
the exhumation history of the orogen. For example,
Mesalles et al. (2014) took advantage of the topographic
relief along the east flank of Mt Beidawu, 3091 m (Fig. 1),
and constructed an age-elevation profile using partiallyreset and totally reset ZrnFT ages. The profile showsan acceleration in exhumation cooling at about 3.2 Ma.
A more recent, integrative study that included new andpreviously published ZrnFT and ZrnHe data, including
Mesalles et al. (2014) shows an additional acceleration at~1.5 Ma (Lee et al. 2022a).
Two periods of accelerated exhumation are also consistent with fission track ages of totally reset detrital apatite from the southern Backbone Range. Available AFTages, which have a lower annealing temperature thanzircon for fission tracks, form two clusters generally distributed on the east and west flanks of the range (Fuller
et al. 2006), excluding three older AFT ages of 27.9, 5.5,
and 4.0 Ma from the southern tip where erosion has beenminimal. Along the eastern flank and the Hengchun Peninsula, AFT ages range from 0.4 to 2.4 Ma, whereas along
the western flank, AFT ages are slightly older, ranging
from 3.6 to 2.4 Ma. The older ages on the western flankoccur in the hanging wall of the Chaochou Fault (Fig. 2,
Site L) (Shyu and Sieh 2005; Chen et al. 2005; Yang 1986)
and cluster along the map trace of the fault. The fault isan east-dipping thrust and appears to record substantial
displacement as mildly metamorphosed slates have beenthrust west over Holocene alluvium of the Pingtung Plain(Liu et al. 2009). The limited age range and the distribution of samples in the hanging wall of the thrust suggest
the Chaochou was active at about 3.6–2.4 Ma, whichcoincides with the time of accelerated exhumation cooling in the rocks east of the fault (Mesalles et al. 2014).
The cluster of younger apatite ages along the eastern
flank may reflect the acceleration in exhumation at ~1.5

Page 11 of 50

Myr proposed by Lee et al. (2022b), which is based primarily on ZrnHe ages of detrital zircon.

4 Methods
4.1 Field data and structural analyses

Our field studies focused on the outcrops previously
mapped as the Yuli Belt. However, we also completed
reconnaissance excursions in the adjacent Tailuko Belt
and the sedimentary cover sequences associated withthese belts. Our goals in the Tailuko Belt were to confirm
the presence and geometry of the youngest foliation, typically “S3” in this belt, identify additional late-stage structures associated with the ductile to brittle transition, andcollect samples of vein materials for fluid inclusion anddating analyses.
Structural data and samples were collected primarily
during four field campaigns in January 2018, July 2018,
January 2020, and April 2021. The January 2018 trip
was focused on the Dalun River, near its junction with
the Xinwulu River. The campaign in July 2018 sampled
seven river sites, encompassing over 100 km along thetrend of the orogen (Fig. 1), whereas the January 2020
trip was a reconnaissance trip along short stretches of theSouth Cross-Island Highway, the Heping River, and the
Kukutsu Cliffs (Fig. 1). The April 2021 campaign included
an eight-day campaign along the lower reaches of the
Chingsuei and reconnaissance trips to the Taimali and
Zhiben Rivers in southern Taiwan. Fieldwork included
making detailed observations of mesoscopic structuresand lithologies, collecting structural data (e.g., foliations and lineations), and collecting oriented samples for
microscopic analysis.
Field sites were along river systems due to the thickvegetation cover and rugged topographic relief. The
island straddles the Tropic of Cancer, and the summer
monsoons produce a thriving and diverse biosphere anda thick jungle cover. However, river channels are often
sinuous and produce polished exposures of the metamorphic basement rock with both along-strike and crossstrike views. River sites studied here are shown in Figs. 1
and 2.
Correlating a penetrative fabric or deformation event
across different lithostratigraphic and tectonostratigraphic units is also challenging (Tobisch and Paterson
1988). These challenges are particularly acute in highly
sheared rocks, e.g., S–C mylonites and phyllonites, wheremultiple foliations can be generated during a single
deformation event. The late-stage veins, however, are acommon and relatively conspicuous feature in essentially
all units studied and, in most cases, led to the identification of additional late-stage structures. At some outcrops, when the number of earlier foliations could not beeasily determined, the late-stage structures were taggedByrne et al. Progress in Earth and Planetary Science(2024) 11:23

preliminarily as part of Dy (for ­Deformationyoungest), with
Sy representing the youngest foliation and Ly representing the youngest lineation. We then integrated previous work with the new observations and developed adeformation history for each unit to provide context for
the late-stage structures. In most cases, the deformation stages and fabric classifications developed by previous workers were maintained, although this resulted
in inconsistent subscripts for the youngest fabric acrossdifferent units. To resolve these inconsistencies, we propose an accretionary prism model where more deeply
buried rocks that may record multiple deformational“events” are juxtaposed next to rocks that are accreted atshallower structural levels that record only the youngest
deformational event. We also propose a new deformational event and fabric that appears limited to the Tailuko Belt and the western part of the Yuli Belt near the
Tailuko.
Additional details on sample collecting and field andpetrographic analyses are provided in Additional file 1:
Appendix SI.
4.2 Fluid inclusion microthermometry

Samples used for fluid inclusion microthermometry(FIM) were sourced from three localities representative of the metamorphic basement (red squares, Fig. 1).
The northern site is in the Tailuko Belt and is exposed
along the Kukutzu Cliffs, Hualien County. The central
site is located in the Yuli Belt along the lower reaches of
the Shoufeng River. The southern site was situated along
the Xinwulu River in the eastern Tailuko Belt, near thecontact with the Yuli Belt. Specimens were late-stage
veins or single euhedral crystals 1–4 cm in length. Doubly polished thin sections of four quartz crystals andone adularia crystal were produced for petrographic andmicrothermometric analysis.
A detailed description of the methods used for fluidinclusion microthermometry, including preparation of
doubly polished thick sections, is in Additional file 2:
Appendix SII.
Once constructed, the thick sections were photographed at low magnification (0.5–4X) and examined for
usable fluid inclusions. The fluid inclusion section is thenbroken into small pieces required for microthermometricmeasurements, as described by Touret (2001). A delineation of the broken fragments as well as other thermometric data can be visualized on the fluid inclusion map. The
map is required to reconstruct the broken fragments intheir original position for later SEM-CL analysis. Withineach individual chip, fluid inclusion assemblages wereclassified as primary, secondary, or pseudosecondary.
An ideal area of interest was chosen to collect the microthermometric data based on the number and clarity ofPage 12 of 50

fluid inclusions in the field of view. Detailed sketches
were produced depicting individual fluid inclusions
within the area of interest (Shepherd et al. 1985; Touret2001). Before the heating and cooling experiments were
performed, observations of the fluid inclusions wererecorded at standard temperature and pressure: the sizeand shape of the inclusion, the number of type of phasespresent, and the L:V ratio, percent vapor bubble, or
degree of fill (Shepherd et al. 1985).
40
4.3 
Ar/39Ar dating40

Ar/39Ar analyses were performed at the University ofVermont Noble Gas Geochronology Laboratory. Mineral separates were loaded into aluminum foil packets,arranged in a suprasil vial, and placed in an aluminum
canister for irradiation. Samples were irradiated at the
Oregon State University Radiation Center in the CLOCIT
facility for 18 h with multigrain aliquots of Fish Canyon
Tuff Sanidine to act as a flux monitor (28.201 ± 0.046 Ma;
Kuiper et al. 2008). Laser step heating for 40Ar/39Ar dating was conducted with a Santa Cruz Laser Microfurnace 75W diode laser system. Three aliquots of adularia
from sample T19HC-01, one crystal each, were loaded
directly into wells in a copper sample holder. The gas
released during heating was purified with SAES getters,
and argon isotopes were analyzed on a Nu Instruments
Noblesse magnetic sector noble gas mass spectrometerin peak-hopping mode during step-heating analyses.
Data from samples and flux monitors were corrected forblanks, mass discrimination, atmospheric argon, neutron-induced interfering isotopes, and the decay of 37Ar
and 39Ar. Correction factors used to account for interfering nuclear reactions for the irradiated samples are from
Rutte et al. (2018). Decay constants are from Min et al.
(2000) and Stoenner et al. (1965). Mass discrimination
was calculated by analyzing known aliquots of atmospheric argon for which the measured 40Ar/36Ar value wascompared with an assumed atmospheric value of 298.56(Lee et al. 2006). A linear interpolation was used to calculate J factors for samples based on sample position
between flux monitor packets in the irradiation tube. The
data analyses were achieved using both an in-house data
reduction program and Isoplot 3.0 (Ludwig 2003).
4.4 Strain

We have taken multiple approaches to understanding
the strain and strain histories of the youngest deformation, although the difficult access to outcrops limited systematic analysis. At the broadest scale, the orientation of
the youngest penetrative fabric and associated late-stage
veins consistently show vertical shortening and subhorizontal extension throughout the metamorphic core.
The high angle between veins and the youngest foliation

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23(~90°), and the limited dispersion of the vein orientations
also argue for vertical thinning and a low degree of noncoaxial strain (kinematic vorticity < 1.0). Observations ofstructures in sections parallel to the principal strain axes
(i.e., X, Y and Z) provide additional constraints on the
kinematics. For example, structural data from a deformed
pebble conglomerate along the Mugua River and an
extensive exposure of phyllonite along the Chingsuei
River provide a preliminary deformation history for the
Yuli Belt and, therefore, context for the late-stage structures. Finally, we integrate these new results with a previously published study of the finite and incremental strain
histories in the Yuli Belt sedimentary cover, mapped as
the “Eastern Slate Belt” (Mondro et al. 2017).

5 Results
Structural mapping of the metamorphic core focused
primarily on transects across the Yuli Belt, althoughreconnaissance studies were also completed in partsof the Tailuko Belt and the lower-grade sedimentary
sequences interpreted to represent higher structural levels of the metamorphic rocks. We start with field studies of the lower-grade sedimentary rocks, one from thearea of Zhiben River and the second along the eastern
Chingsuei River. Four additional but more limited studies focus on the higher-grade rocks exposed along the
middle and western reaches of the Chingsuei River andalong short sections of the Hongye and Mugua Rivers(Fig. 1). We also re-examined rocks along the Dalun Riverthat expose the Shoufeng fault and rocks along Heping
River and Kukutus cliffs where late-stage structures havebeen dated isotopically. Structural data collected during more limited studies along the Fuyuan, Matian, andShoufeng Rivers are also included in our compilation.
Despite the differences in their structural complexity, theunifying theme of these areas is the apparent progression from west-northwest-directed thrusting and penetrative stretching to northeast extension. Integrationof our studies with the available literature suggests thatthe late northeast extension event is characterized by (1)a low-dipping pressure solution cleavage, (2) mineralfilled veins that formed late in the development of thepressure solution cleavage, and (3) a weak to moderately
developed stretching lineation that is generally perpendicular to the veins. The general aspects of these characteristic structures are discussed after the site-specificdescriptions.
5.1 Zhiben River

Observations along the Zhiben River, south of exposures of the Yuli Belt (Fig. 1), reveal a relatively mildly
deformed sequence of sandstones and shales. Sandstonelayers often grade upward from coarse-grained sands

Page 13 of 50

to fine, laminated, or convoluted layers, consistent withturbidite deposition. Soft-sediment deformation consisting of slump-like structures, similar style folds, suggesting high ductility, debris deposits that include sections ofoverturned bedding (e.g., Fig. 3), and isolated fold hinges
are also relatively common. The soft-sediment structures
are cross-cut by a weak to moderately developed pressure
solution cleavage that dips steeply to gently east (Fig. 4).
Only one penetrative foliation, ­S1, was recognized inthis area, consistent with the more regional study by
Conand et al. (2020) (Fig. 3C). At one locality along theriver’s north bank, sandstone layers form a northwestverging asymmetric fold with the pressure solution cleavage forming an axial planar foliation. Graded bedding(Fig. 3A), however, shows the layers are upside-downwith layering in the short limb facing east-southeast
rather than west-northwest in the direction of fold vergence. One interpretation is that the outcrop represents
the overturned limb of a larger regional-scale fold. The
fold asymmetry and cleavage dip suggest a downwardfacing, east-dipping, overturned antiform. Alternatively,
overturning of the sedimentary units may have occurred
before folding and cleavage development. In fact, Pelletier and Hu (1984), Lu et al. (2001) and Chang et al.
(2007) have mapped overturned sedimentary units of asimilar scale in the area of the Taimali River and argued
they had a sedimentary origin. Our observations are consistent with this interpretation.
Kinematics of the asymmetric folds exposed along theZhiben River can be inferred from the trend of associated fold axes, the dip direction of the axial planar cleavage, and the trend of fibrous strain fringes around rigid
objects (probably diagenetic pyrite) (Figs. 3, 4). The foldaxes suggest movement to the northwest, whereas thecleavage generally dips east, suggesting a more westerly
motion. Strain fringes exposed along the plane of cleavage trend east-northeast to west-southwest, oblique tothe movement directions inferred from the dip of cleavage and the fold axes. Conand et al. (2020) also recognized a similar pattern of northeast-trending stretching
oblique to the trend of measured fold axes, with a significant portion of the data showing overlapping trends(Fig. 3B, C). They proposed that folding and stretchingdeveloped contemporaneously in a left-lateral wrenching environment during oblique convergence (see also Lu
and Lee 2001).
5.2 Eastern Chingsuei River and highway 30

Further north, along the north bank of the Chingsuei
River and along Highway 30, near the junction with the
Lakulaku River (Fig. 5), historical outcrops of the Miocene Chulai Formation initially exposed two decimeterscale west-to-northwest-verging asymmetry folds similarByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 14 of 50

C)
A)

N
0
33

30

60

S0

0

E

W

W

E

S1

Bo
tto
m

12

0

24

0

To
p
0

21

0

15

S

B) Poles to S0; open squares:
fold axes; red dot cylindricalbest t axis (19 205)
Fig. 3 Outcrop photograph and orientation data from sedimentary rocks along the Zhiben River interpreted to represent higher structural levelsof the Yuli Belt. A Photograph of cleavage, S1, dipping east and overturned bedding, S0, dipping west. B Poles to bedding (n = 16), cylindrical best fitof bedding, and fold axes (n = 2). C Fold axes from Domain II of Conand et al. (2020)

to the fold along the Zhiben River. Unfortunately, thefold along Highway 30 that showed clear asymmetry has
been covered with concrete and is no longer available forobservation. Graded bedding in the inferred short limbs
of both folds also shows(ed) younging to the east rather
than to the west (Fig. 6), suggesting that the beds were
overturned before folding, consistent with observations
along the Zhiben River. The top-to-the-west shear is also
consistent with outcrop observations in the Chulai Formation to the west (see below). However, the fold asymmetry of the preserved fold along the Chingsuei River(Fig. 6) is less clear than for the concrete-covered fold,allowing alternative interpretations. For example, if thebeds were not overturned before folding, the east-facing
bedding on the Zhiben River may reflect clockwise rotation associated with top-to-the-east shear.
A few 100 m west of the fold described in Fig. 6, a lowdipping phyllonite is exposed and S–C fabrics confirm
the top-to-the-west motion inferred from the asymmetric folds (Fig. 7). The “C” surfaces dip gently east andare approximately parallel to lithologic layering, making it impossible to identify primary sedimentary layering in this area. The “S” surfaces also dip east but at asteeper angle, indicating a top-to-the-west sense of shear(Fig. 7B, C). The S surfaces are locally associated withwest-verging folds of C surfaces or “P” shears consistent
with west vergence. Multiple S and C surfaces, P shears,
and locally extensional shear bands result in an irregularor wavey fabric at an outcrop scale.
The decimeter-scale west-verging folds and the S–C
phyllonites in these two areas are cross-cut by a moreplanar foliation that dips gently northeast (Fig. 8) andappears undeformed by the S–C fabrics. The crosscutting relations, however, are more conspicuous in thesteeply dipping limb of the fold where lithologic layering is at a high angle to the younger fabric. One of thelate foliation surfaces also exposes an elliptically shaped
reduction halo with its long axis parallel to a weak mineral lineation in the cleavage plane (Fig. 8). Relatively
undeformed to mildly deformed quartz veins are alsopresent in this area of the Chingsuei River and shownortheast extension (Fig. 8) parallel to the long axis of thereaction halo and mineral lineation.
The decimeter-scale fold and the phyllonites pre-datethe low-dipping foliation and primary sedimentary structures in the steep limb of the fold suggests little deformation occurred before folding. We, therefore, interpretthe fold and S–C phyllonites as D1 structures and theByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 15 of 50

B)
A)South view

E)
N

30

0

0

30

33

33

N
30

12

0

24

E
12

0

0

24

15
0

0
15

21
0

0

W

0

E

21

W

60

0

60

0

S

S

C) S1 cleavage planes and poles;
Blue: stretching lineationD) Late-stage extensional veins

Fig. 4 Photographs of S1 in sedimentary rocks along the Zhiben River. A and B are approximately parallel to the XZ plane of the strain ellipsoid
inferred from strain fringes around pyrite framboids. Framboid diameters average about 0.5 cm. Enlarged framboid and fringe B suggest finite
stretch (1 + e1) of 3.5. C Orientation data for S1 (planes and poles) stretching lineation from strain fringes. D Orientation data for late-stage veins(planes and poles). E Orientation data for stretching lineations from Conand et al. (2020)

low-dipping foliation and late-stage veins as representing
D2. This interpretation is consistent with Lin et al. (1984),
who proposed that the Chulai Formation recorded two
deformation events.

5.3 Middle and western Chingsuei River sites

Further west, more limited studies along the middle andwestern sections of the Chingsuei River (Fig. 5) reveal
excellent exposures of zones of relatively high strain in
the Yuli Belt (see, e.g., photographs in Figs. 9, 11). RocksByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 16 of 50

uR

lak
ku

La

Yuli
Be

X

Tailu
k

Chu
lai
Fm

o Be

lt

lt

.

X

Western
Chingsuei

9, 10Middle Chingsuei

7

X
6, 8

X = Hwy 30
outcrops

Eastern
Chingsuei

11, 12= Locations of serpentinite bodies and smaller
blocks of pillowed greenstone, epidote
amphibolite, plagiogranite and glaucophane
schist.
Fig. 5 Location of focus areas along the Chingsuei River. See Fig. 1 for location and text for discussion. Numbers refer to Figure locations; “X”
to outcrops along Hwy 30

along the middle Chingsuei River comprise alternating layers of quartz-rich and quartz-poor metamorphosed sediments, typical of the Luntienshan memberof the Yuli Belt (Lin et al. 1984). However, primary sedimentary structures are completely obliterated, and theoutcrops are often dominated by highly strained sedimentary rocks with lenses and stringers of quartz thatprobably represent early quartz veining. Most of thelenses of quartz have a symmetric shape with respect tothe main foliation. However, isolated hinges of isoclinal
folds are not uncommon, and a few hinges are suggestive of asymmetric folds (Fig. 9). The isolated hinges areinterpreted to be remanent of an earlier foliation, making the main foliation in the outcrop ­S2. One ellipticallens of quartz is oriented at an oblique angle to the ­S2,
suggesting a clockwise rotation, viewed to the north, anda top-to-the-west sense of shear (Fig. 9C). ­S2 in theseoutcrops is also deformed by a set asymmetric folds andan associated spaced crenulation cleavage (Fig. 9). The

axial surfaces and cleavage dip moderately east and thefold asymmetries consistently indicate a top-to-the-west
sense of shear. Based on cross-cutting relations and relations across Fig. 10A (e.g., western Chingsuei River) andalong strike (e.g., Dalun River), we interpret this younger
crenulation cleavage to represent S3.
Exposures along the western Chingsuei River comprise relatively thick quartz- and feldspar-rich metamorphic rocks, including a Mn-rich garnet schist often
found in the Chingsuei Planation member of the Yuli
Belt (Lin et al. 1984). Like the exposures along the middle Chingsuei River, the units are intensely deformed,
with quartz-rich rocks approaching mylonites (Fig. 11A)
and quartz-poor rocks displaying phyllonitic textures.
However, the main foliation dips steeply to moderately
west, in contrast to the foliation in the middle Chingsuei
River (Fig. 11B). As in the middle Chingsuei, S–C shear
structures suggest a top-to-the-west sense of shear. The
main foliation is also deformed by a set of asymmetricByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 17 of 50

A)
S0

W

E

S2

B)
N
30

0
33

30

60

0

p
To

W

E
m

tto

Bo

12

0

24

0
15

21
0

0

S

C) Bedding, S0
Fig. 6 Outcrop photographs and orientation data for S0, and S2 from the Chulai Fm along the north bank of the eastern Chingsuei River (see Fig. 5
for location). A Photograph of a recumbent, east-facing syncline and axial planar foliation, interpreted to be S2. East-facing bedding may reflect CW
rotation of upright beds during top-to-the-east folding or CCW rotation of overturned beds during top-to-the-west folding. B Inset of A shows
bedding facing east. C Orientation of bedding and interpreted fold axis. Curved arrow shows one interpretation for sense of rotation during folding.
See text for discussion
folds that typically have relatively gentle to open interlimb angles. In some cases, the interlimb angle decreases
substantially as the amplitude of the fold increases, producing tight to isoclinal folds and suggesting progressive shear approximately parallel to the axial surface(Fig. 11A). The folds consistently verge to the west (i.e.,
west-side-down) and, based on cross-cutting relations
and observations a few km along strike (i.e., Dalun River,discussed below), we interpret this later deformation to­D3 and label associated folds and associated foliations, F3and S3, respectively.
The eastern, middle and western sections of the Chingsuei also display differences in the dip direction of thedominant foliation, typically lithologic layering which
is subparallel to “C”. Along the eastern Chingsuei thedominant fabric dips northeast (Fig. 7B), whereas along
the middle Chingsuei, the dominant fabric dips northgently (Fig. 10A), and along the western Chingsuei, theByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 18 of 50

A)

View north

S1c
S1

s

S1c

S1c

N

N
30

60

0

30

0

0

30

33

33

30

60

0

EW

0

0

15
0

0
15

21

12

0

24

0

12
0

0

24

E

21

W

S

S

B) S-C relations: Red: S1s,
black: S1cC) S1 cleavage poles, S and C
undi erentiated; Red: F1 fold axes.
red square: cylindrical best t

Fig. 7 Outcrop photograph of S–C phyllonite in the Chulai Fm along the north bank of the eastern Chingsuei River. This location is west of Fig. 6
and may represent a transition from the Chulai Fm to the Yuli Belt. See text for discussion. Sense of shear inferred from the S–C mylonite fabricshown with two opposing arrows. B S–C relations with curved arrows showing sense of rotation. C ­S1 and S–C planes, undifferentiateddominant fabric dips west or northwest steeply (Fig. 11B).
This regional-scale change in dip direction across strikesuggests a regional-scale antiform. The middle and western sections also crop out on the east and west sides,
respectively, of relatively large bodies of serpentinite andsmaller blocks of pillowed greenstone, epidote amphibolite, plagiogranite, and glaucophane schist (Fig. 5),
consistent with an antiform. East–west cross sections
constructed by Lin et al. (1984) and Lin (1999a) also show

an antiformal structure centered on the middle and western sections of the Chingsuei. A similar north–northeasttrending antiform occurs along strike in the Yuli Belt (Yiet al. 2012) (Fig. 2 site F). However, the data from thewestern Chingsuei site is constrained by only two foliation data points (Fig. 11B). Without facing indicators,
we are also not able distinguish between a regional-scale
antiform or an overturned, east-facing synform between
the middle and western Chingsuei sites. The exposures

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 19 of 50

A)

Y

X

North view.
N

N
30

60

E
12

0

24

12

0

24

0

0

0

21

E

15

15

0

0

W

0

W

60

0

21

0

30

0

0

30

33

33

30

S

S

B) Poles to S2 (black), F2 axial
surfaces (red), and stretching
lineation (blue)C) Planes of late-stage veins

Fig. 8 A Outcrop photograph of a stepped exposure that provides an oblique view (lower) and profile view (upper) of S2. The oblique view
includes a reduction halo on S2. Long and intermediate axes suggest a finite stretch (1 + e1) of 3.0. B Orientation data for S2, F2 axial surfaces, and L2.
C late-stage veins; brown arrows show interpreted kinematic axes based on observed slip-along veinsalong the western Chingsuei near the trace of theShoufeng fault may have also been rotated CW (viewed
north) as proposed by Crespi et al. (1996) for rocks along
the SXIH and CXIH.

Finally, the penetrative foliations exposed along boththe middle and western sites, ­S2 and ­S3, are cross-cut
by mildly deformed quartz and calcite-filled veins that
dip steeply southwest (Figs. 10C, 12C) and are slightly

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 20 of 50

A)

S3
F3

F2

F3

F3

F2

North view.

S3
FIg 9B

B)

F3

S3

S3
S3

S2

S3

C)

S2

F3

Fig. 9 Outcrop photographs of S–C phyllonite in the Yuli Belt along the north bank of the middle Chingsuei River (see Fig. 5 for location). Main
foliation, ­S2, dips gently north (Fig. 10A and is horizontal in the photographs. Isolated isoclinal fold hinges with one suggesting an S-shaped
asymmetry (black arrows) indicate a pre-existing foliation. Elliptical quartz inclusion oblique to S2 (C, inset) suggests a top-to-the-west senseof shear during S2. D3 fold axial surfaces, F3, and crenulation cleavage, S3, dip east (Fig. 10A). Scale, shear sense and view direction are the same in all
photographs

wrinkled or folded by a moderately developed foliationthat dips gently northeast (Figs. 10B, 12C). Along thewestern Chingsuei, strain fringes around euhedral pyrite
crystals are parallel to a weak mineral lineation on thelate foliation and appear related to the development of

the late, low-dipping foliation (Fig. 12B). The fringes andmineral lineation record northeast stretching parallelto the extension direction inferred from the late quartz
veins (Fig. 12C). Based on cross-cutting relations andobservations along strike (i.e., Dalun River, discussed

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 21 of 50

N
30

0
33

30

60

0

W

E
12

0

0

0
15

21
0

24

S
A) S2 foliation planes (black)and S3 foliation planes (red)
N

N
30

60

0

30

0

0

30

33

33

30

E W

0

0

0

0

0

15

15

21

12

0

24

0

12

0

24

E

21

W

60

0

S

S

B) S4 planesC) Late-stage veins

Fig. 10 Structural orientation data from the S–C phyllonites in the Yuli Belt along north bank of the middle Chingsuei River. A Orientation of S2and S3. Late-stage foliation, S4, and associated late-stage veins dip gently north and steeply southwest, respectively (B and C)

below), we interpret the late-stage veins and associatedfoliation to represent D4.
5.4 Dalun and Xinwulu Rivers

The Dalun River hosts a series of meanders subparallel to
the Shoufeng Fault (Stanley et al. 1981) and joins the Xinwulu River near the eastern end of the SXIH (Fig. 13A).
In this area, the Tailuko comprises a marble-rich unit
composed of various quartz and mica schists, chert, anddiscontinuous white and light gray marble layers. Outcrops along the Dalun and closest to the Xinwulu showa moderately west-dipping mylonitic foliation that Stanley et al. (1981) and Ho et al. (2022) labeled S2, based oncross-cutting relations and regional patterns. Ho et al.
(2022) also recognized a sub-horizontal stretching lineation and well-developed asymmetric folds and lenses ofquartz and marble, consistent with left-lateral shearing.
Our mapping along the Dalun River south of the junction with the Xinwulu shows ­S2 having a moderate tosteep northwest dip (Fig. 13B), consistent with Stanley
et al. (1981) and Ho et al. (2022). We also mapped a moderately west-dipping foliation (Fig. 13B) that deforms S2

Byrne et al. Progress in Earth and Planetary ScienceA)(2024) 11:23

View southwest

S2

S3
N
0

30

33

30

60

0

W

E
12

0

24

0

21

15

0

0

S

B) S2 (black) and S3 (red)Fig. 11 Outcrop photograph and structural data of phyllonite
in the Yuli Belt along south bank of the western Chingsuei River(see Fig. 5 for location). A Photograph of well-developed phyllonite,
parallel to pencil and later shear zone interpreted to be S3. View
southwest. B Orientation data for S2, and S3, fabrics in A. Movement
direction (black arrows) is inferred from asymmetric folds associatedwith S3 and shows west-side down

and is axially planar to west-verging asymmetric folds of
S2 This foliation is equivalent to S3 of Stanley et al. (1981)
and may have been included as part of D2 in Ho et al.
(2022). The S3-related folds typically have relatively gentle to open interlimb angles. However, in many cases, theinterlimb angle decreases substantially as the amplitudeof the fold increases, producing tight to isoclinal folds
and suggesting progressive shear approximately parallel to S3. Although these structures were not studied inPage 22 of 50

detail, we interpret the asymmetric folds and progressive
shear zones as D3 “P” shears (Fossen 2010) related to thewestward thrusting of the Yuli above the Tailuko rather
than part of D2.
Finally, exposures along the Dalun and Xinwulu Rivers
also show a younger, gently dipping foliation (Fig. 13D)that appears to have formed contemporaneously with thedevelopment of late-stage quartz and calcite veins that
dip steeply northeast (Fig. 13C). This late fabric appears
undeformed by S3 and we have labeled it S4. However,this late fabric is significantly different from S4, as definedby Stanley et al. (1981) and is more similar to Stanley’s
F6 event, which includes a weak, gently dipping cleavage
associated with late-stage veins that dip steeply northeast. This gently dipping late-stage foliation is also geometrically and temporally similar to S3 recognized by Ho
et al. (2022) and Ho and Lo (2015) in their studies of the
Tailuko Belt, and S3 recognized by Lin et al. (1984) in theYuli Belt along the Chingsuei River. To reconcile thesedifferent interpretations and maintain consistency acrossunits, we follow Stanley and use S3 for low to moderately
dipping foliation associated with west-verging asymmetric folds along the Dalun River and S4 for the moregently dipping and younger foliation associated with latestage steeply dipping veins. That is, we recognize S3, as
defined by Stanley et al. (1981), as a penetrative fabricthat formed in association with west-directed thrustingof the Yuli Belt above the Tailuko Belt. We apply a similarinterpretation to the asymmetric folds along the western
Chingsuei (see above and Fig. 11A) near the map traceof the Shoufeng fault and suggest that similar interpretations may apply to other structures near the Shoufeng.
However, some areas may also preserve a transition fromdominantly strike-slip, D2, to dominantly west-directedthrusting, D3.
5.5 Hongye River

Exposures along the Hongye River are more limited,
although one outcrop near an abandoned outwash control structure exposes the core of an overturned synform(Fig. 14) cut by a gently dipping foliation associated withsteeply dipping late-stage quartz veins. The fold and associated minor folds trend northeast and displays an axial
surface that dips moderately southeast (Fig. 14A, B), consistent with early formed folds in the Yuli Belt. The foldgeometry suggests that its part of a northwest-verging
asymmetric fold, again similar to other folds in the Yuli
Belt.
The late-stage structures that cross-cut the outcropscale fold are interpreted to represent D4 structures andinclude a penetrative foliation, S4, lineation, L4, and shear
bands (Fig. 15). A suite of relatively late veins, generally perpendicular to stretching lineation on S4, are also

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 23 of 50

A)

View southwest
S4

B)

View northwest

N
0

30

33

30

60

0

W
12

0

0

24

15
0

21
0

C) S4 planes (black), latestage veins (dashed) andstretching lineation (red)E

S
Fig. 12 Outcrop photographs and orientation data along the south bank of the western Chingsuei River (see Fig. 5 for location). A Two late-stage
quartz veins and F4 axial surface (red dashed line). B Quartz-mica schists with quartz-filled strain fringes around euhedral pyrite. C Orientation data
for S4 cleavage, L4 stretching lineation, from the strain fringes, and late-stage veinspresent (Fig. 15A, C). The geometry of S4 and the shear
bands suggest they represent either extensional shears orplanes of penetrative shortening. The former would indicate a top-to-the-north sense of shear, whereas the latter
would indicate the opposite sense of shear.

Thin-section observations in X–Z sections of ­S4 provide additional information in resolving the kinematics.
For example, albite porphyroblasts record an interior
foliation observed by linear arrays of opaque minerals, including graphite (Fig. 16). The interior foliationis characterized by gentle to open folds, but the host

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 24 of 50

N
0

30

33

A)

Xinw

ulu

R.

30

60

.
Ch
u

Dalu

lai

Yu
li

Be

Fo
rm

lt

Ta
ilu
ko
B

elt

0

nR

.

st
ru
th
e
t
d)
e
a
ic n ebr zo turn
r
Im
e
v
(o

W

E
12

0

24

= Structural data0

1 mile

0

21

0

15

S

B) S2 (black) and S3 and S4,undi erentiated (red)
N

30

60

0

30

0

30

33

0

30

33

N

W

60

0

E

W

E

Vx

12

0

0

0

15

0

0

15

21

12

0

24

0

0

24

21

VxVx

S

S

C) Late-stage veins
D) S4 planes and poles
Fig. 13 Geology and structural orientation data around the Dalun and Xinwulu Rivers junction. A Google image and geologic map showing
an imbricate thrust zone between the Tailuko and Yuli Belts (from Stanley et al. 1981) and data sites for this study. B Orientation data of S2 (black
great circles) and S3 and S4 undifferentiated (red great circles). Orientation data for late-stage veins (C) and S4 (D)albite grains display no sign of crystal plastic deformation. The trends of the interior and exterior foliations are also different by as much as 50° and indicate
a top-to-the-southwest asymmetry. Strain shadows are
infilled with chlorite, albite, and, less commonly, muscovite. Ilmenite porphyroblasts at a high angle to thefoliation yield relatively large strain shadows, resulting
in chlorite strain fringes with a low L:W ratio. Strings offine-grained phyllosilicates in the fringes estimate theinstantaneous stretching axis and indicate a top-to-thesouthwest sense of shear. We therefore interpret theshear bands observed in the outcrop (Fig. 15A) as compressional structures (i.e., “S” surfaces) associated with
top-to-the south-southwest shear (Fig. 15A, C).

5.6 Mugua River

Along the north bank of the Mugua River, in the northern exposures of the Yuli Belt (Fig. 1), a stretched pebble
conglomerate provides additional insights into S3 and S4(Fig. 17). The outcrop exposes a meter-thick shale-matrix
conglomerate interlayered with sandstone, phyllite, andgreenschist, preserving at least in part a stratigraphic succession. The units, however, are intensely deformed, andthe pebbles, composed primarily of chert, appear flattened and highly elongated. The long axes of the pebbles
parallel lithologic layering, the dominant foliation in thearea, and plunge gently northwest (Fig. 17B). The shapesof the deformed clasts measured on oriented slabs yield
similar values for Y/Z and X/Y, suggesting plane strain.

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 25 of 50

A)

NE view, parallel to L3 & FA1
N
0

30

33

30

60

0

W

E
12

0

24

0

21

15

0

0

S

B) Lithologic layer (bedding?)
and early-stage fold axes
Fig. 14 Photograph and structural data of an outcrop-scale synform in the Yuli Belt along the Hongye River. A Synform fold axis trends northeast
parallel to the change in slope where the person is sitting and axial surface dips moderately southeast (right). B Orientation of lithologic layering(bedding?) (black great circles) and fold axes of smaller folds associated with the synform

The primary foliation dips gently north and is locallycross-cut by a crenulation cleavage that dips gently northeast (Fig. 18). The intersection lineation alsoplunges gently northeast, approximately perpendicular
to the long axis of the pebbles. The two foliations andthe stretching direction provided by the elongated pebbles suggest an S–C relation. At this point, however, wehave not confirmed whether the crenulation cleavage
is “S,” which indicates top-to-the-northwest, or C’ anextensional crenulation cleavage, which would indicate
top-to-the-southeast.
We also recognized a low-dipping foliation interpreted to be S4, and a suite of late-stage veins (Fig. 19A,
B) that appear unaffected by the S–C foliations. The
veins are generally orthogonal to the low-dipping foliation, and to a weak mineral lineation on S4, suggesting
they are part of the late-stage structures recognizedthroughout the eastern Backbone Range.

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 26 of 50

A)

qtz)

qtz)

B)
SE view

N

N
30

60

0

30

0

0

30

33

33

30

W

60

0

E W

E

Y
12

X

0

12

0

0

24

0

0

0

21
0

15

15

21
0

24

S

S

C) S4s (black), S4C (blue) and L4(red). X and Y show view directionof photographs in Fig 16

D) D4 extensional veins

Fig. 15 Photograph of outcrops and orientation data associated with phyllonite along the gently dipping limb of the synform in Fig. 14. A Outcrop
photograph of rock face approximately parallel to the XZ plane of youngest foliation, S4. Inset B shows stretching lineation on the youngest foliation.
Lineation trends southwest (C), approximately perpendicular to late-stage extensional veins in A and plotted in D. S–C fabrics, blue and black
great circles in C, respectively, and rotated porphyroblasts (Fig. 16e, f ) indicate a top-to-the south-southwest sense of shear. Based on correlations
with fabrics observed in the western Chingsuei River the late fabrics are interpreted to be D4 structures

5.7 Heping and Kukutsu areas

We completed reconnaissance fieldwork in the areasof the Heping River and Kukutsu Cliffs (Fig. 1) to confirm structural patterns identified by previous workers
and collect additional structural data and samples for40

Ar/39Ar dating. The Heping River area exposes outcrops of the Tachoshui gneiss whereas the Kukutsu Cliffs
exposed the Kanagan gneiss (Lin and Chen 2016). Bothsuites of rocks are interpreted to Cretaceous intrusions
and display a well-developed north-northwest dippingByrne et al. Progress in Earth and Planetary Science(2024) 11:23

mylonitic foliation with an associated mineral stretching
lineation that plunges gently west (Ho et al. 2022; Korren
et al. 2017; Chen et al. 2022; Jiao 1991). Ho et al. (2022)
consider the mylonite to be the second foliation, S2, in the
Tailuko Belt, and we follow their terminology here. Wanget al. (1998) obtained ages ranging from 4.1 to 3.0 Ma
for the mylonite based on 40Ar/39Ar crystallization agesof biotite. The mylonite is cross-cut by pseudotachylitelined normal faults (Korren et al. 2017) that yield 1.5 Maages using 40Ar/39Ar (Chen et al. 2017a), consistent withthe older ages from the mylonite, which it cross-cuts.
The mylonites are also cut by north-striking, verticalveins and a penetrative but sparsely developed foliationthat dips gently southwest (Figs. 20). The late-stage veinsdisplay a variety of thicknesses and contain quartz, calcite, chlorite, and, less commonly, adularia or sanidine,
epidote, and phengite (Fig. 20A). A weak mineral stretching lineation, L­ 3, associated with the low-dipping foliation plunges moderately east-northeast, parallel to theextension direction inferred from the late veins (Fig. 20B,
C). A late-stage extension shear zone less than a centimeter in thickness and associated with the late-stage
foliation also yielded an east-side down sense of shear(Fig. 20D). These stretching lineations are also subparallel to striations associated with the formation of thepseudotachylite (Fig. 2, site B) (Korren et al. 2017). We
interpret these late-stage lineations to be ­D4 structures.
Chen et al. (2022); (see also Wang 1998) obtained 1.1 and
1.5 Ma ages for two samples of adularia from the veins inthis area based on 40Ar/39Ar, and additional isotopic agesof adularia from these outcrops are discussed below. Following Chen et al. (2022), we infer that the extensional
veins and pseudotachylite formed contemporaneously
and propose that D4 also formed at this time.
Finally, the Heping and Kukutsu areas are in the northern limb of the Backbone Range orocline, and late-stage
structures appear to have rotated ~48° CW viewed down
about a vertical axis relative to similar structures south ofPage 27 of 50

the orocline (Fig. 1). This rotation is consistent with GPS
data from the RYUK block, which shows the northernlimb of the orocline rotating at ~47°/Myr, suggesting thatthe orocline was initiated ~1 Ma. In our compilation oflate-stage structures (Fig. 21), we show the observed orientation of the structures from Heping and Kukutsu andthe orientation of the structures after 48° of rotation.
5.8 General characteristics of the late‑stage foliation, ­S4

On the regional scale, the youngest penetrative foliationin the eastern Backbone Range dips shallowly to moderately northwest or southeast, and a compilation of datacollected along the eastern Backbone Range (Fig. 21) suggests a gently northeast plunging antiform. These results
are consistent with mapping in the central part of the Yuli
Belt that shows a broad, elongated antiform centered on
the Yuli Belt (Gungfu area Yi et al. 2012) (Fig. 2) and witha limited number of more detailed studies in the Tailuko,
Yuli and Slate Belts (Fig. 2) (Ho and Lo 2015; Lee 1995;
Clark et al. 1992; Stanley et al. 1981) that suggest the
youngest foliation forms one or more gently northeast
plunging antiforms. The youngest foliation is also associated with a suite of late-stage extensional veins that consistently record northeast extension (Fig. 21). Northeast
extension during the development of the youngest foliation is also consistent with limited observations of thefinite strain. A geometrically and kinematically similar
fabric also occurs in the lower-grade rocks that crop out
south and east of the Yuli Belt and may represent shallow
structural levels of a Late Miocene accretionary prismassociated with the Yuli Belt (Chen et al. 2019). Finally,a systematic study of the kinematics associated with theyoungest foliation suggests a slight bias toward a top-tothe-southwest sense of shear.
At an outcrop scale, the late-stage foliation ranges
from a low-angle penetrative pressure solution cleavage
in lower-grade units to a low-angle crenulation cleavage
often parallel to axial surfaces of open-to-isoclinal folds.

(See figure on next page.)
Fig. 16 Photomicrographs of structures produced by D4 deformation in the Yuli Belt along the Hongye River. Schematic in bottom left of eachphotograph illustrates the interaction of S4 (red) with previous foliations or other structures. White arrow shows the coordinate direction,and plunge and trend indicate the view direction and pole to the plane of the thin section. See Fig. 15C for orientation data. a, b (XPL) show
albite-rich (top) and albite-poor (bottom) domains defining the cleavage. c Quartz grains within microlithons that show grain shape preferredorientation, slight undulose extinction, and grain boundaries indicative of sub-grain rotation and grain boundary migration recrystallization. d
Relationship between a relatively early formed lineation, defined by an aggregate shape preferred orientation (labeled L3-1, ASPO and a laterlineation defined by a grain shape preferred orientation (labeled “L3, GSPO”). The early formed lineation is interpreted to represent D­ 2, whereas
the later lineation is interpreted as L4. Photograph is parallel to XY plane of S4. e Albite porphyroblasts within a cleavage domain depicting
an interior foliation and over-growth zoning. Acicular rutile with ilmenite coatings, tourmaline, and apatite inclusions are common, especiallyin grains with a deformed interior foliation. Locally, one coated ilmenite grain is isoclinally folded, while another depicts extension. f SEM-BSE
image of image in e, depicting the variation in composition. Albite porphyroblasts contain ilmenite inclusions and are surrounded by chlorite
and muscovite strain fringes. Different orientations of internal and external foliations indicate a top-to-the-right sense of shear, which in geographic
coordinates is top-to-the-southwest. Orientation data are shown in Fig. 15

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 28 of 50

Fig. 16 (See legend on previous page.)

The folds can be inclined or recumbent depending on
the initial orientation of the folded foliation (Fig. 22B).
For example, in the Tailuko Belt, where the older foliations typically dip steeply to moderately west, D4 folds
are relatively common and are generally asymmetric withshort limbs that appear to have rotated counterclockwise(when viewed to the north). Fold axes also plunge shallowly to the northwest (Ho et al. 2022), consistent withsuperposition of the low-dipping S4 on steeply to moderately dipping layering. Folds are less common in the Yuli
Belt, where the older foliations have more variable dips.
The common occurrence of quartz- or mica-rich schists

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 29 of 50

A)

0

33

30

N
30

60

0

W

E
12

0

24

21

0
15

0

0

S
B) S2 planes, poles, andpebble long axes (blue)(n=9).
Fig. 17 Outcrop photograph (A) and orientation data (B) of a matrix-supported pebble conglomerate in the Yuli Belt along the north bankof the Mugua River

also results in a rough to dispersed cleavage rather than aplanar pressure solution cleavage, making identifying S4
more challenging.
In thin section, S4 is defined by late-stage, ductile
deformation, including continuous foliations, S–C
foliations, extensional crenulation cleavages, and strain

fringes around relatively rigid inclusions, including
albite, pyrite, and ilmenite porphyroblasts (Fig. 16). S4
is observable by ductile/brittle deformation and a retrograde mineral assemblage, including quartz + albite +muscovite + chlorite + others that often fill S4-related
strain fringes. Cleavage domains and high-strain zones

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 30 of 50

A)

S2c

S2

s

S2c

S2s

View ~northeast
N
0

30

33

30

60

0

W

E

12

0

0

0
15

21
0

24

S

B) S-C relations: Red: S2s,
black: S2c and pebble longaxes, blue
Fig. 18 Outcrop photograph and orientation data of intercalated mudstone and greenschist capped by greenstone and interlayered
with the pebble mudstone shown in Fig. 17. A Photograph of rock surface approximately parallel to XZ plane of S2 strain ellipsoid. Interpreted S
and C surfaces associated with ­S2 are labeled. 1-inch drill hole for scale. B S–C relations for individual outcrop. Curved arrow: rotation axis interpreted
from S–C intersection; black arrow: movement direction inferred to be perpendicular to S–C intersections and shows top-to-the-northwest. Blue
dots show trend and plunge of pebble long axes (from Fig. 17)

contain muscovite, chlorite, albite, and graphite. SynS4, mineral-filled, mode-I fractures are also common
where S3 is developed.

5.9 General characteristics of the late‑stage veins

The late-stage, mineral-filled, mode-I fractures, orveins, track the distribution of the youngest penetrative(2024) 11:23

Page 31 of 50

N
0

30

0
33

33

N

30

30

60

W

E
12

0

24

0

0

W

E
12

0

24

0

0

0
15

15

21

60

0

21
0

0

30

Byrne et al. Progress in Earth and Planetary ScienceSSA) S4 planes and poles (n=5), andL4 stretching lineation, red (n=3)B) Late-stage veins: planes andpoles (n=13)

Fig. 19 Orientation data from the Yuli Belt along the north bank of the Mugua River (Fig. 18): A Late-stage foliation and associated stretching
lineation and B quartz veins. Based on correlations with fabrics observed in the western Chingsuei River, the late fabrics are interpreted to be D4
structures

foliation through the core and cover rocks. The veins consistently dip moderately to steeply southwest (Fig. 21),
except in the northernmost sites where the Backbone
Range has been rotated clockwise (Fig. 1). The veins generally range in thickness from 0.25 to 20 mm but canbe up to 10–15 cm and are often filled or partially filledwith quartz + albite + calcite +/− adularia. Rare mineral inclusions of muscovite and ilmenite were observed.
In outcrop, the veins appear generally undeformed andcross-cut the major fabrics. However, closer inspection(e.g., hand lens and petrographic observations) reveals
that they show slight ductile deformation and are locallyoffset by the penetrative foliation (Figs. 22, 23E, F).
In thin sections, the veins show evidence of brittle/ductile strain. The center of the fracture-fill displays euhedral to subhedral, equigranular arrays of quartz, albite,and adularia that show uncommon undulose extinction
and grain boundaries indicative of bulging recrystallization and less common sub-grain rotation recrystallization. These grains also display elongate-blocky to fibrousvein mineral morphologies (Fig. 23A–D); (see, e.g., Bons
2000). Elongate grain morphologies perpendicular to
the edge of the vein formed by the crack-seal mechanism (Ramsay 1980) and indicate the veins formed as
repeated mode-I fractures as fluid pressure exceeded thetensile strength within the rocks (Secor 1965). The elongate crystals are generally quartz and albite, 50–200 µm
in length, and optically continuous to each other. Theyshow little to no undulose extinction, and inclusion trails

parallel to the vein edge are common. The mode-I fractures most commonly display syntaxial growth morphologies, indicating the most recent growth was within thecenter of the vein.
5.10 Fluid inclusion analysis of late‑stage vein materials

Fluid inclusion microthermometry on quartz and adularia crystals from the veins was also utilized to provide
constraints on the pressure and temperature conditions,and estimates of the metamorphic fluid compositions
during vein formation across the metamorphic core. Previous fluid inclusion microthermometry studies in thecore are uncommon. On the westernmost edge of the
Slate Belt, along the Lishan fault, amorphous quartz veinswith two-phase fluid inclusions record homogenization temperatures between 208 and 225 °C and salinitiesof 4.9% NaCl (Chan et al. 2005). The continued mineralization of the quartz, as well as deeply sourced albite(sodium), potassium feldspar, and rare earth elements atcooler, epithermal conditions, was attributed to a steep
fluid overpressure gradient maintained by the overlying
slate cap-rock. On the eastern edge of the Slate Belt, near
the contact with the Tailuko Belt and the CXIH, Crawet al. (2009) used fluid inclusion microthermometry on
silicic, ankeritic, and chloritic veins from the Pingfengshang gold mine. Entrapment temperatures ranged from
240 to 350 °C at inferred depths ranging from < 1 km to
10 km. The workers combined temperature data withthermochronological data from Beyssac et al. (2007) toByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 32 of 50

A)

N
30

0
33

30

60

0

W

E
12

0

24

0

21
0

0
15

S

South viewB) Late-stage vein planes; redtriangle shows mean pole
N

N
30

60

E
12

0

24

0

0

E

S4
Ex shear zone
12

0

24

0

15

0

0

15

21

W

0

W

60

0

21

0

30

0

0

30

33

33

30

S

S

C) S4 and L4,
blue dotD) S4 and associate
extensional shear zone

Fig. 20 Outcrop photograph and structural orientation data from Kukutsu Cliffs (Fig. 1 for location). A Photograph showing euhedral adularia
crystals approximately 0.5 cm in width growing along the wall of a late-stage extensional vein. B and C show orientation data for late-stage
veins and youngest penetrative foliation, respectively. D Orientation data for S4 and an associated extensional shear zone that indicate
a top-to-the-east-southeast sense of shear. Based on correlations with fabrics observed in the western Chingsuei River, the late fabrics areinterpreted to be D4 structures. The Kukutsu Cliffs outcrops are located on the RYUK block (Fig. 1), which is rotating clockwise with respect
to the central Backbone Range (Rau et al. 2008). See text for discussion

generate P–T–t paths and determine exhumation ratesof 2 mm/yr for the rocks in the western slate belt. The
orientation of late veins associated with conjugate kinkbands at Pingfengshang are consistent with the orientation of the late-stage veins discussed here.
Our fluid inclusion microthermometry data (seealso tables and fluid inclusion maps in Additional files
3, 4, 5, 6, 7, 8, 9, 10, 11, 12, 13, 14) indicate maximum

temperatures during vein precipitation of between 200
and 450 °C. These temperatures are consistent with
deformation textures of contain quartz and feldspar
grains in cleavage microlithons that display undulose
extinction and evidence of sub-grain rotation, bulging,and grain boundary migration recrystallization. Thesegrain boundaries are consistent with deformation fromapproximately 350–500 °C (Fossen and Cavalcante 2017;

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 33 of 50

Vein Orientation Data Northof the Orocline

LATE-STAGE VEIN AND FOLIATIONORIENTATION DATA
= Vein Planes= Vein Poles

N

N

33

N=19
n=3

0

W

12

0

E

0

0

21

0
21

S

N

15
0

30

33

30

0

30

33

30

0

S

S

N

0

15

0

0

24

24

N=35
n=9

15

x.
pro
ap line
c
oro

12
0

12

0

0

E

21
0

24

W

n=11

E

60

0

Matian R.

60

0

W

30

= Stretching
lineation= # S4

30

60

0

N
30

= # veins

30

30

0

0

30

Shoufong R.

= Foliation Plane, Mean

n

33

33

= Vein Poles, MeanN

After CCW

Observed

Mugua R.
60

0

N=12
n=38

60

0

W
W

E

E
12

0

24

0

0

21

0

0

15

12

0

24

S
0

21

0

15

S

30

0

N
0

30

33

N=20
n=6

N
33

Hongye R.

30

30
0

Fuyuan R.

60

W
W

E

E
12

0

24

0

12

0
24

0

21

15

0

0

0

21

0

15

S

S

Xinwulu /Dalun R.

N
0

30

33

N
0

30

33

N=3
n=163

N=13
n=9

60

0

00

30
0

Chingsuei R.
N=18
n=5

60

60

W
W

E

E
12
0

0

24
12

0

24

0

21

15

0

0

0

21

0

15

S

S

N
0

30

33

30

Zhiben R.
N=11
n=13

60

0

W

E
12

0

24

0

0

21

0

15

S

Fig. 21 Synthesis of new late-stage structural data from nine sites along the eastern Backbone Range. Data includes late-stage veins (planes, poles
and mean poles), foliations and stretching lineations; not all sites have all types of data. Chingsuei River site is compilation of data along middle
and western reaches. Northern most site shows orientation data for Kukutsu Cliffs and Heping River before and after rotation associatedwith the Backbone Range orocline. Arrows show extension direction inferred from vein orientations, stretching lineations and strain fringes

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 34 of 50

Fig. 22 Outcrop photographs and photomicrographs of veins from the Yuli Belt. White arrow shows the coordinate direction
for photomicrographs; plunge and trend indicate the view direction and pole to the plane of the thin-section photomicrographs. Schematicin bottom left of each photograph illustrates the interaction of S4 (red) with previous foliations or other structures. a Late-stage veins dipping
south and cross-cutting S4. b Late-stage vein cross-cutting S2 and S4 characterized by slight ductile deformation along S4 planes. c XPL imageof a late-stage vein at a high angle to S4, featuring elongate, optically continuous albite grains. d SEM-BSE image of the late-stage vein in c. The veincontains quartz, adularia, and ilmenite. e SEM-BSE image of a late-stage vein containing quartz, adularia, and ilmenite featuring elongate to blocky
grain morphologies. f SEM-BSE false color image of D2/D4 fabric, ­D4 microstructures defined by the retrograde paragenesis in Ti-oxide strain
shadows and of a late-stage vein at a high angle to the fabric. The vein contains quartz and chloriteByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 35 of 50

Fig. 23 Photomicrographs of veins from the Yuli Belt. White arrow shows the coordinate direction, and plunge and trend indicate the view
direction and pole to the plane of the thin section. a, b (XPL)) Late-stage vein cross-cutting the S4 fabric at a high angle. The vein features quartz,
adularia, and ilmenite. Elongate grain morphologies are observable along the vein/wall rock contact, and blocky grains with slight undulose
extinction are dominate in the center of the vein. c, d (XPL) Late-stage vein composed of quartz, albite, and uncommon chlorite cross-cutting S4
characterized by ductile deformation along an S4 cleavage domain. Elongate grain morphologies are observable along the vein/wall rock contact,and blocky grains with slight undulose extinction are dominant in the center of the vein. e, f (XPL) Thin vein with adjacent grains showing a grain
preferred orientation indicating northeast-side-down normal offset

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Passchier and Trouw 2005). Within the late-stage veins,evidence of grain boundary migration recrystallizationis less common and may suggest slightly cooler temperatures during strain accumulation from the ­S4 fabric(~< 400 °C).
The regional fluid system is composed of ­H2O–CO2–
NaCl. High ­XCO2, low ­XCO2, and aqueous inclusions
commonly occur in similar fluid inclusion assemblages
or at least within the same sample chip. Homogenization temperatures range from 199 to 350 °C, and salinities range from 0.4 to 9.4 mass% NaCl. The variations
in homogenization temperatures and salinities outlined
below are attributed to the lithology or relative structural
position of the country rock at each of the three sites.
Euhedral quartz and adularia crystals from the Kukutsu
Cliffs, the northern site, preserve two-phase (LV, liquidand vapor), ­H2O–NaCl inclusions ranging from < 5 to
50 µm in size, with an average of ~10 µm (Fig. 24a, b).
Shapes were amorphous or negative crystal-polygonal insome adularia inclusions (Fig. 24A, B). Microthermometry was performed on twenty-eight fluid inclusions
between two samples. Homogenization temperatures
range from 199 to 229 °C +/− 5 °C with an average of
214 °C +/− 5 °C. Ice melting temperatures range from− 2 to − 0.2 °C +/− 3 °C with an average of − 0.87 °C +/−3 °C. These melting temperatures correspond to salinities
ranging from 0.4 to 3.4% mass NaCl (Fig. 25). The quartz
and adularia samples from the northern site depict nocompositional heterogeneity in SEM-CL. This suggeststhat all observed fluid inclusions are likely primary or
pseudosecondary.
Quartz crystals from the central site along theShoufeng River host primary and secondary, three-phase(LaqLCO2V) inclusions ranging from < 5 to 120 µm in size,with an average of ~40 µm; shapes are generally amorphous (Fig. 24c, d). Microthermometry was performedon eighty-eight fluid inclusions between two samples.
Homogenization temperatures range from 240 to 350 °C+/− 5 °C with an average of 310 °C +/− 5 °C. The average ­CO2 homogenization temperature was 30.1 °C +/−5 °C. The melting temperature for the solid ­CO2 phasewas ~− 57.5 °C +/− 3 °C. Clathrate melting temperatures range from 4.7 to 9.4 °C +/− 3 °C with an averageof 8.7 °C +/− 3 °C. Calculated salinities range from 0.4 to
9.4%mass NaCl (Fig. 25). The central samples depict rare
growth zoning and common, post-formational, structural, and chemical heterogeneities observed as linear
arrays of lower intensities in the CL image. Large, isolated inclusions and inclusions associated with growthzoning or no post-formational heterogeneities wereinterpreted as primary or pseudosecondary if along acompositionally homogeneous microcrack. Fluid inclusions associated with an assemblage parallel to linear

Page 36 of 50

heterogeneities in cathodoluminescence images wereclassified as secondary.
One quartz crystal from the southern site along the
Xinwulu River hosts primary and secondary, two-phase(LV) inclusions ranging from < 5 to 75 µm in size, withan average of ~40 µm; shapes are generally amorphous(Fig. 24e, f ). Microthermometry was performed on sixty
fluid inclusions. Homogenization temperatures range
from 282 to 340 °C +/− 5 °C with an average of 305 °C+/− 5 °C. Ice melting temperatures ranged from − 5.4 to− 1.5 °C +/− 3 °C. Calculated salinities range from 2.6
to 8.5%mass NaCl (Fig. 25). Distinct from the northernand central samples, the southern samples displayed conspicuous growth zoning parallel to crystal boundaries. A
linear array of compositional heterogeneity was observed
by cathodoluminescence in regions where growth zoning
was less prevalent. The fracture-like arrays are < 500 µm
in width and cross-cut growth zoning. Fluid inclusions
in assemblages parallel to crystal boundaries or growthzoning observed by SEM-CL were classified as primaryor pseudosecondary. Fluid inclusions within the crosscutting, lower-intensity fracture pattern were classified as
secondary.
The low magnitude and range of homogenization
temperatures suggest that the D­ 4 fracture system was
exposed to metamorphic to meteoric fluid along a hydrostatic to hydrothermal geothermal gradient. The rangeof pressure at constant temperatures for ­CO2 inclusions(Fig. 26) suggests growth during protracted coolingunder typical hydrostatic conditions with thermal gradients ranging from 30 to 80 °C/km. Salinities range from
0.4 to 9.4%, suggesting local mixing of surface fluids andfluids rising from the zone of regional metamorphism.
Quartz crystals may locally contain chlorite habits or rareinterstitial chlorite, which may have ions sourced fromthe zone of greenschist facies metamorphism connectedby fluid advection. The depth of vein precipitation ranges
from < 1 to 7 km (Fig. 26), which supports a dynamic,
tectonic-hydrothermal fluid system and constrains thetransition to brittle extension at a depth of ≤ 7 km.
40
5.11 
Ar/39Ar analysis of adularia

Two samples of adularia were collected for 40Ar/39Ar
dating from coastal exposures along the Kukutzu Cliffs(Fig. 1). The samples consisted of mm to cm-sized euhedral crystals collected by hand from late-stage veins in
the Kanagan gneiss (Lin and Chen 2016). In one of thesamples, however, the crystals were coated with microchlorite and appeared dark green. The second sample
lacked the chlorite coating and displayed a pearl-whitecolor, which is more typical of the adularia from theseoutcrops (Fig. 20A) (Huang 1953). Perhaps not surprisingly, the chlorite-coated sample yielded a complex

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 37 of 50

Fig. 24 Photomicrographs of fluid inclusion types found in quartz and adularia crystals from late-stage, D4, veins from northern (a and b), central (cand d), and southern sites (e and f). Primary, two-phase H­ 2O-NaCl inclusions in quartz (a) and adularia (b) from the Kukutsu that display a negative
crystal form. c and d show primary, three-phase H­ 2O–CO2–NaCl inclusions in quartz from the central, Shoufeng River, site. One inclusion in c
was locally decrepitated at 271 °C. Fluid inclusions in d show a planar fluid inclusion assemblage H­ 2O-CO2–NaCl. e Linear fluid inclusion assemblage
characterized by pseudosecondary three-phase H­ 2O–CO2–NaCl inclusions from the southern, Xinwulu River, site. In f, also from the Xinwulu Riversite, planar fluid inclusion assemblages are characterized by pseudosecondary and sub-contemporaneous secondary two-phase, ­H2O–NaCl
inclusions

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 38 of 50

Fig. 25 Bar plot depicting the distribution of homogenization temperatures per sample, including ­H2O–NaCl, and ­H2O–CO2 type inclusions,and a scatter plot depicting the weight% of NaCl for the bulk fluid inclusion as determined from the ice melting temperatures shown in the upperbar graph. The left inset map of Taiwan depicts the sample localities (see also red boxes Fig. 1)

degassing history and an inverse isochron with large
errors and geologically unreasonable ages. The late-stage
veins from which the adularia was sourced are oriented
357°/69° and S4 dips 48° toward 334° in this area. Mineral paragenesis for the vein fill was commonly quartz +albite + calcite +/− adularia. Microstructural observations and CL imaging of the adularia suggest little postformation deformation or strain accumulation within the
crystal structure (Fig. 22). Sample data tables, step-heating experiments, and inverse isochrons for T19HC-03 areincluded in Additional file 15: Appendix SIII. All errors
are quoted at the one sigma level.
Sample T19HC-01, which lacked the thin chlorite
coating, yielded a mean age from three aliquots of
2.38 ± 0.15 (1σ) Ma (Fig. 27). Three single-grain aliquots were analyzed by step heating and resulted inrelatively well-defined inverse isochrons with isochron
statistically indistinguishable ages of 2.54 ± 0.13 Ma,
2.24 ± 0.13 Ma., and 2.37 ± 0.14 Ma with MSWDs =2, 105, 7. Apparent age spectra from step-heating runs
show older apparent ages within the first 10% of 39Ar
released. Intermediate stages show initially anomalously old step(s) followed by a loss profile. Minimum
ages associated with the intermediate temperature steps
that variably define argon-loss profiles are indiscernible

but generally < 1 Ma. Although discordant, one aliquot
yielded a weighted mean age of ca 2.3 Ma.
In conjunction with the geologic setting, the stepheating runs suggest that both samples, T19HC-01 and
T19HC-03 were affected by partial resetting possibly
associated with hot metamorphic fluids propagatedthrough the fractures. Such an influx is also supportedby the high geothermal gradients suggested by FIM onquartz and adularia and the lack of notable solid solutions within the adularia as suggested by CL analysis
and the sub-concordant nature of the higher temperature steps for most of the runs.
Chen et al. (2022) also dated two samples of adularia
from the area of the Kukutsu Cliffs using 40Ar/39Ar andobtained slightly younger inverse isochron ages of 1.09
± 0.03 Ma and 1.5 ± 0.3 Ma. The available isotopic ages,
therefore, show that horizontal extension occurredfrom c. 2.4–1.1 Ma.
5.12 Kinematics associated with the late‑stage structures

A review of outcrop and thin-section observations suggests a low degree of non-coaxial strain (kinematic vorticity < 1.0) during ­D4. At all field sites, late-stage veinsare consistently at a high angle to ­S4 and ­L4, suggestinga thinning strain regime and the absence of a significant

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 39 of 50

Fig. 26 Scatter plot depicting the minimum entrapment temperature and calculated pressure conditions for ­H2O–CO2–NaCl fluid inclusions.
The linear plots are lines of constant density (isochores) for the average density plus/minus one standard deviation for the ­H2O–NaCl inclusions
with the density listed in g/cm3. The dashed lines are hydrostatic (H) and lithostatic (L) thermobaric gradients for thermal gradients between 30
and 80 °C/km. The shaded areas define the statistical range (~1σ) of P–T conditions during which the ­H2O–NaCl fluids were trapped. The peaktemperature of the Yuli Belt is 550 °C (Kouketsu et al. 2019) based on RSCM, which we assume was the maximum entrapment temperature
for the fluid inclusions. The maximum entrapment temperature for the southern site was constrained at 450 °C using RSCM (Beyssac et al. 2007). The
depth was calculated using an overlying rock density of 2.7 g/cm3

component of rotational strain (Ring et al. 1999; Feehan
and Brandon 1999; Fossen and Cavalcante 2017). In addition, observations parallel to ­S4 and ­L4 show a sense ofshear indicators that are generally weakly to moderately
developed, but a top-to-the-south sense of shear is dominant (e.g., Fig. 16e, f ). For example, 12 of 17 thin sections
preserve well-defined indicators, with 9 recording topto-the southwest, 2 recording top-to-the northeast, and 1
showing top-to-the northwest (Fig. 28). The sedimentary
sequences south of the Yuli also appear to preserve a consistent top-to-the-south sense of shear (Lu and Lee 2001;
Conand et al. 2020). The top-to-the-south-southwest
sense of shear is consistent with asymmetric structuresassociated with the pseudotachylite in the northern Tailuko Belt but inconsistent with the sense of shear interpreted for the ductile shear zones in the Slate Belt northof the pseudotachylite (Fig. 2B), after rotations related tothe orocline are removed.

6 Discussion
The late-stage veins, cleavages, and ductile shear zones
that define the youngest penetrative deformation along
the length of the eastern flank of the Backbone Range
consistently show vertical shortening and north-tonortheast extension, suggesting that tectonic exhumation
played a role in exhuming and cooling the metamorphic
core. To evaluate this interpretation, we begin with a discussion of the age of late-stage structures and proposethat the structures post-date the formation of S2 and S3
in the Tailuko Belt and pre-date the formation of the orocline in the northern Backbone Range. We then consider
the transition from D2, recorded by strike-slip deformation primarily in the Tailuko Belt, and the late-stage
structures, delineated as D4. We propose that this transition marks the westward extrusion of the high-pressure
rocks mapped in the Yuli Belt. Extrusion was apparently
accommodated by thrust faulting in the west, recognizedByrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 40 of 50

Fig. 27 Ar step-heating and inverse isochrones for three aliquots of sample T19HC-01. See text for discussion. Sample data tables, step-heating
experiments, and inverse isochrons for T19HC-03 are included in Additional file 15: Appendix SIII

extensional shear zone remains cryptic, the change fromhorizontal to vertical shortening throughout the metamorphic core argues that tectonic processes played an
important role in the exhumation of the high-pressure
rocks in the Yuli Belt.
6.1 Age of the youngest penetrative structures

Fig. 28 Mean S4, L4 and shear sense associated with S4/L4
in XY planes at the river sites in Fig. 21. 12 of 17 samples
yielded well-defined indicators, with 9 showing top-to-the-SW,
2 showing top-to-the-NE, and 1 showing top-to-the-NW.
A top-to-the-southwest sense of shear is consistent with observations
by Lu et al. (2001) and Conand et al. (2020) in the southern Backbone
Range

as D3, and displacement along an extensional shear zone
in the east evidenced primarily by a low-dipping foliationand steeply dipping extensional veins, D4. Although the

The relative age of the youngest penetrative fabrics in theeastern Backbone Range varies across different tectonic
units, although the fabrics have similar geometries andkinematics. For example, the sedimentary rocks south of
the Yuli Belt, including the exposures along the Zhiben
River and Domain II of Conand et al. (2020) preserve asingle penetrative foliation and lineation, S­ 1 and ­L1, and asuite of late-stage mineral-filled veins. Further north andeast of the Yuli exposures, along the eastern section of
the Chingsuei River, and along a short stretch of Hwy 30,
the sedimentary sequences preserve two foliations that
we interpret as S1 and S2. Mondro et al. (2017), mapping
more regionally but in correlative rocks (i.e., the Chulai
and Hsinkao Formations), also documented two penetrative events. The two events are recorded by two sets offibrous strain fringes, with the early formed set recording generally west-northwest to east-southeast stretching
and the later set recording east-northeast to west-southwest stretching. We propose that these stretching events

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

correlate with S1 and S2 documented along the eastern
Chingsuei River. Further west in the higher-grade metamorphic rocks of the Yuli and Tailuko Belts, S3 is alsopresent, and the youngest foliation, S4, dips gently northand is associated with northeast stretching.
Based primarily on the geometry and kinematics ofthe structural fabrics and the absence of detailed stratigraphic control, we propose an accretionary prism model
where more deeply buried rocks are more deformed.
During exhumation, units with more protracted deformation histories can be juxtaposed next to less deformed
rocks, and both suites may then experience a relatively
young deformation event that “stitches” the units. In thiscontext, S4 in the Yuli and Tailuko Belts is the “stitching”
event recorded as ­S2 in the Chulai and Hsinkao Formations and as S1 in the low-grade rocks south of the Yuli.
Isotopic ages and cross-cutting relations of dated
events constrain the absolute age of the youngest structures. For example, the gently dipping foliation, extensional shear zones, and veins consistently cross-cut the S2
mylonite in the Tailuko Belt, which ranges in age from 4.0
to 3.1 Ma based on 40Ar/39Ar dating (Wang et al. 1998).
We infer, therefore, that 3.1 Ma represents the maximum
age of the youngest structures. Several of the structures
included in the suite of youngest structures, e.g., a lowangle pressure solution fabric, extensional veins, andshear zones, have also been rotated by the orocline innorthern Taiwan, which appears to have started forming ~1.0 Ma., providing a minimum age for the youngest
structures. This range for the maximum and minimum
age is consistent with three 40Ar/39Ar ages of adularia collected from the late-stage veins, 2.4 +/−, Ma (this study),
1.5 Ma +/− and 1.1 +/− Ma (Chen et al. 2022), with a40
Ar/39Ar determined age of pseudotachylite of 1.6 Mathat fills a suite of extensional fault zones (Chen et al.
2017a) and 40Ar/39Ar ages of mica from the Daguan shear
zone, ~0.7 Ma which cross-cuts S­ 3 (redefined here as­S4) in the Yuli Belt (Yi et al. 2015). We therefore proposethat the suite of relatively young structures in the eastern
Backbone Range formed between ~2.4 and 0.7 Ma, with
deformation possibly starting as early as ~3.1 Ma.
6.2 Exhumation history

A full understanding of the exhumation of the high-pressure rocks in the Yuli Belt from their estimated depth of~50 km is beyond the scope of this paper. However, we
recognize three processes that may have driven overthickening in the rear of the wedge and led to the transition to structural thinning implied by geometry and kinematics of the late-stage structures: (1) changes in local
buoyancy by chemical and phase changes (e.g., hydrated
blocks of mantle), (2) crustal-scale pressure gradients ina subduction channel during continuous subduction, andPage 41 of 50(3) changes to the thickness and composition of the subducting crust (Fossen 2010; Warren 2013). Our preference is for (3), although (1) and (2) may have also played
a role. We favor (3) primarily because regional compilations of low-temperature geochronologic data from themetamorphic core show the formation of a major outof-sequence thrust, the Tayuling thrust, at ~6 Ma (Fig. 2
site M)(Lee et al. 2022b), and the thrust carries the Tailuko Belt in its hanging wall, arguing for the incorporation of Eurasian continental crust (i.e., SE China) at thistime (Lan and Liou 1981; Dewangga et al. 2021; Leeet al. 2022b). This interpretation is also consistent withthe recent documentation of Taiwan-derived sediments
arriving in the foreland basin in the Late Miocene to
Early Pliocene, ~5–4 Ma (Hsieh et al. 2023). We, therefore, propose that subduction of thick continental crust(e.g., equivalent to the Tailuko Belt) ~6–5 Ma initiatedexhumation of the high-pressure rocks (Yuli Belt). Continued exhumation is documented with isotopic ages ofretrograde metamorphism, including a 4.4 ± 0.1 Ma ageof phlogopite using 40Ar/39Ar, (Lo and Yui 1996) and a 3.3± 1.7 Ma age for zircon in a clinozoisite vein in nephrite
using U/Pb (Yui et al. 2014). This documented decreasein timing is followed by the initiation of structural thinning at ca. 2.4 Ma.
Fluid inclusion microthermometry, vein textures, and40
Ar/39Ar age analyses on adularia collected from thelate-stage veins provide additional information on structural thinning during the ductile to brittle transition.
For example, grain morphologies within the veins indicate they developed under an extensional stress regime
between < 1 and 7 km, with a maximum depth for mineral precipitant ranging from 4 to 7 km. The decrease ofpressure at constant temperatures, as recorded by FIM,
suggests continued brittle/ductile deformation as therock package was exhumed through the hydrothermal
fluid system. Once the overburden was reduced enoughto allow the fractures to remain open, well-developedeuhedral crystals grew in open voids, or the veins werefilled with equant blocky grain morphologies, probably
at depths between 0 and 2 km depth. This transition is
also suggested by the change from crack-seal vein material near the walls of the veins to euhedral crystals in thevein centers. Secondary fluid inclusions at similar trapping conditions were likely formed within the crystals
during the protracted cooling. The cessation of the fluid
system is represented by the deposition of soluble ions in
the form of chlorite, ankerite, calcite, and gold deposits(Craw et al. 2009). If we assume that the adularia crystals
formed before the transition to lithostatic to hydrostatic
conditions (e.g., 2–7 km) at 2.38 Ma, tectonic thinningand erosion rates of between 0.8 and 3 km/Myr would
be needed to remove the overburden. Using the younger

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

40

Ar/39Ar age of ~1.1 Ma for adularia growth suggests
significantly faster rates of 1.8–6.3 km/Myr.
6.3 Extrusion and changing structural regimes

An important but often overlooked consequence of tectonic extrusion in a subduction zone is that the horizontal motions of the extruding blocks can add to theshortening associated with plate convergence (Chemenda et al. 1995; Warren et al. 2008). This additional
shortening component could be accommodated in thefold-and-thrust belt and even equal the extension in therear of the wedge (Chemenda et al. 1995; Warren et al.
2008). In Taiwan, tectonic exhumation of the Yuli Belt
driven, for example, by pressure gradients in the subduction channel, could have accelerated slip along the decollement beneath the accretionary wedge west of the Yuli
Belt, essentially pushing the wedge westward.
To better understand how this process may apply to
Taiwan, we start with a reconstruction of the tectonic
setting near the end of D2 in the Tailuko Belt (Fig. 29A).
We have removed the oroclines in northern and southern Taiwan and consider the orogen a relatively straightsystem consistent with regional-scale strike-slip (Ho et al.
2022). During D2, highly oblique plate convergence wasaccommodated through deformation partitioning (Fitch
1972; McCaffrey 1992) with orogen-normal shortening in
the Slate Belt and orogen-parallel strike-slip in the Tailuko Belt (Ho et al. 2022). For the D2 strike-slip regime,
we envision a relatively simple instantaneous strainstate where the minimum stretch, S3, trends northwest,
parallel to the plate convergence vector, the intermediate stretch, S2, is vertical, and the maximum stretch, S1,
trends northeast (Fig. 29C).
Near the end of D2, we propose that the westward
extrusion of the Yuli Belt caused a significant change inthe tectonic regime (Fig. 29). In this interpretation, extrusion is accommodated by slip along two east-dipping,regional-scale structures: the Shoufeng thrust that crops

Page 42 of 50

out west of the Yuli Belt and carries the Yuli Belt in its
hanging wall and a more cryptic extensional shear zonethat we propose carries the Luzon forearc in its hanging
wall (Fig. 29B). We also propose that extrusion of the Yuli
Belt offset the pre-existing Tailuko shear zone left-laterally along a generally east–west trending fault, forminga regional-scale releasing bend (Fig. 29B). The Chouchi
Fault in the central part of the Yuli Belt may represent
a similar but less well-developed cross-structure (“G” in
Fig. 2). As oblique convergence continued, the state ofstrain in area of the releasing bend changed from thrusting and thickening (i.e., large vertical stretch) (Fig. 29C,
D3) to vertical thinning and strike-parallel extension(Fig. 29C, D4) . That is, formation of the releasing bend
generated an area of low vertical stretch that accommodated vertical extrusion of the Yuli Belt.
In support of this proposed change in regimes, wenote that slip on the Chaochou thrust (Fig. 2) may haveaccommodated the westward displacement associatedwith extrusion of the Yuli Belt. For example, the unusual
north–south strike of the Chaochou Fault (see Figs. 1, 2)
may reflect progressively lower slip, or slip rates, from
north to south along the fault. The presence of a regionalscale, late-stage (i.e., post-cleavage) antiform in thenorthern extent of the Chaochou hanging wall (Huang
and Byrne 2014) is consistent with larger slip along thefault in the north. Progressively lower slip north to south
also correlates with the narrowing of exposures of theYuli Belt in the core of the orogen consistent with thenotion that westward extrusion of the Yuli Belt added tothe slip along the Chaochou. Finally, cooling ages of reset
detrital apatite grains from the hanging wall of the fault
range from 3.6 to 2.4 Ma, which overlap with exhumation
cooling ages in the Yuli Belt (Lo and Yui 1996) and slightly
pre-date initiation of northeast extension recorded by thelate-stage veins that date from 2.4 to 1.1 Ma (this paper
and Chen et al. 2022).

(See figure on next page.)
Fig. 29 Generalized maps, cross sections, and strain states for the Tailuko and Yuli Belts as they are accreted/subducted and extruded. A ~6.0–
3.0 Ma, accretion of Tailuko and Slate Belts and subduction of Yuli Belt as continental crust of Eurasia enters the Luzon subduction zone. Kinematics
and strain partitioning along the plate boundary are from Ho et al. (2022). Schematic cross section X–X’ at the time shown in A. B ~3.0–0.7 Ma,
vertical and westward extrusion of the Yuli Belt accommodated by reactivation of the Tailuko strike-slip shear zone as a high-angle reverse fault,
the Shoufeng shear zone (SSZ). Westward extrusion accelerates slip on the Chaochou Fault (CF in cross section) and offsets the left-lateral shear
zone in the Tailuko Belt, creating a left-stepping releasing bend. Geochronologic data suggests northeast stretching (black arrows) started ca.
2.4 Ma (this study). The location of reconstructed units relative to Taiwan is for general reference. C General strain states for three deformation stages
with “S” representing the stretch (S = 1 + elongation) for each of the principal strains with S1 > S2 > S3. D2: strike-slip deformation in the Tailuko Belt
during oblique convergence and strain partitioning (Ho et al. 2022). D3: east-dipping high-reverse fault between Tailuko and Yuli Belts (represented
by red dashed and continuous lines). Orientations and relative magnitude of stretches are poorly known, but maximum stretch S1 is assumed
to be nearly vertical. D4: vertical thinning and extrusion of the Yuli Belt that minimizes the vertical stretch while the releasing bend during obliqueconvergence maximizes the horizontal stretch. Black dashed line represents late-stage foliation. Note that the maximum stretch trends northeast
for D2 and D4

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Page 43 of 50

Fig. 29 (See legend on previous page.)

Westward displacement of the Tailuko shear zone
as the Yuli Belt was extruded may also explain the limited evidence for left-lateral slip in the area south of thesouthernmost exposures of the Tailuko Belt. For example,

the only detailed geologic map of this area that we areaware of (Usami and Matsumoto 1940) shows two leftlateral strike-slip faults on the east and west sides of the
Tailuko outcrops (Fig. 2 site N). However, map-scaleByrne et al. Progress in Earth and Planetary Science(2024) 11:23

offsets across the faults suggest that they accommodated
relatively small displacements. The eastern fault displaces
an anticline-syncline pair 2–3 km, and this fault is itself
offset 1–2 km by the second left-lateral fault, suggesting a
total slip of ~5 km. Conand et al. (2020) show a regionalscale, left-lateral fault that generally follows the west
boundary of the Tailuko Belt, although only the northernand southern extents have been mapped. Several smaller,
north-striking left-lateral faults have also been mapped inthis area (Lu and Lee 2001; Mesalles 2014), but none ofthese faults have documented offset, and the ages of slip
are generally unknown.
Considering that the Tailuko shear zone is interpreted
to be part of a plate boundary, the total slip is expected to
be much larger than the observed 5 km (Usami and Matsumoto 1940). One option is distributing displacements
related to the Tailuko shear zone throughout the southern backbone range. Alternatively, the westward motionof the Chaochou, which carries the Tailuko Belt in its
hanging wall, raises the possibility that the westward
motion of the extruding Yuli Belt offset the Tailuko shear
zone. This interpretation is attractive because the highlylinear coastline in southeast Taiwan suggests a strike-slip
deformation history. In fact, Zhang et al. (2022), following
Lewis et al. (2015) have recently proposed that the Luzon
Trough, which merges with the Longitudinal Valley in
Taiwan, is, at least in part, a pull-apart basin between two
left-lateral strike-slip faults. The western basin-bounding
fault recognized by Zhang et al. (2022) is equivalent tothe southern extent of the Tailuko shear zone we are proposing here. Westward offset of the Tailuko shear zone is
an attractive interpretation because it would also result ina left-stepping, left-lateral fault system (Fig. 29B), which
predicts northeast-southwest extension in the northeast
quadrant (Fig. 29B), analogous to a pull-apart basin. This
structural setting is consistent with the kinematics of D4
in the Yuli and Tailuko Belts.
In summary, the late-stage structures in the Tailuko andYuli Belts and parts of their sedimentary covers appear toreflect vertical shortening and horizontal extension associated with the extrusion of the Yuli Belt. Vertical shortening reflects the vertical ascent of the extruding mantle,
probably accentuated by a pre-existing, steep strike-slip
fault (see, e.g., Mann and Gordon 1996; Avé Lallemant
and Guth 1990; Platt 1983). Horizontal extension developed as the Yuli Belt was extruded westward up the subduction zone, offsetting the north-striking Tailuko shear
zone left-laterally. This new structural setting createda left-stepping, left-lateral shear zone and a regime ofregional-scale extension in its northeast corner (Fig. 29).
Westward extrusion of the Yuli Belt was accommodated
by additional slip along the Chaochou thrust along thewestern flank of the orogen.

Page 44 of 50

6.4 Structural thinning

The magnitude of tectonic exhumation during the youngest penetrative event can be estimated using the finite
strain values obtained in the lower-grade sequences ofthe metamorphic core (Mondro et al. 2017), although theuncertainties are relatively large. The strain data show amean, generally horizontal stretch of 4, which, assuming constant volume and plane strain, is equivalent toa vertical stretch of 0.25 and indicates vertical shortening of 75%. Assuming a 30 km thick accretionary prism,
a sub-horizontal foliation, and a uniformly distributed
vertical strain rate, vertical shortening would account forover 20 km of tectonic exhumation. The relatively young
event also appears to have been active for ~1.3 Myr (i.e.,from ~2.4 to 1.1 Ma), indicating an exhumation rate of~15.4 km/Myr, which seems unreasonably high.
There are several critical unknowns in estimating tectonic exhumation from the strain data. Ring and Kumerics (2008) and Feehan and Brandon (1999) argue thatstrain values collected from exposed high-pressure terranes typically significantly overestimate the amount ofcrustal thinning. This is because we know very little about
strain rate laws or the distribution of strain at depth, androcks can travel relatively quickly through zones of highstrain rates as they are exhumed. For example, Ring andKumerics, (2008) used a strain rate law that increased
proportionally with depth to estimate the exhumationof high-pressure rocks in the Alboran Sea (exhumed
from 40 km, Platt et al. 1998) and concluded that verticalstretches of 0.03 were required. This relatively high value
implies strain ratios of ~1000, which Ring and Kumerics
considered unrealistic.
Although the vertical stretch observed in Taiwan,
0.25, is substantially lower than the calculated value forhigh-pressure rocks in the Alboran Sea, the estimatedexhumation rates in Taiwan of ~15.4 km/Myr are unreasonable. If we assume vertical shortening started at theend of mylonitization associated with ­S2 in the Tailuko
Belt, or about 3.1 Ma, the tectonic exhumation rate for a
30 km thick crust would still be 10 km/Myr, which is also
relatively high.
One alternative is to assume that vertical shortening
represents strain associated with a regional-scale detachment or shear zone with a finite width (e.g., Fig. 29C, ­D4).
If we assume, for example, that the observed horizontal
stretch value, ~4, represents strain that accumulated as
the rocks moved through the range of pressures and temperatures indicated by the fluid inclusion data, ~5 km, ahorizontal stretch of 4 would remove 3.75 km of overburden and imply an exhumation rate of 2.9 km/Myr. Thisrate is comparable to modern rates of erosion in Taiwan(Dadson et al. 2003; Derrieux et al. 2014).

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

A regional-scale, east-dipping detachment, as shownin Fig. 29C, D4, is attractive as it provides an explanation
for CW rotation (north view) of older fabrics (e.g., reddashed line in Fig. 29B) (Crespi et al. 1996), the CW rotation of the Yuli-Chulai contact (see Fig. 5, Stanley et al.
1981) and the top-to-the-east shear observed the lowergraded metamorphosed sedimentary rocks east andsoutheast of the Yuli Belt (Lu and Lee 2001). Remnantsof the detachment may also be preserved at the base of
the mafic/ultramafic blocks at Wanrong (Fig. 2 between
the Shoufeng and Hongye Rivers) and Ruisui (Fig. 2 at thelatitude of the Hongye River) that are mapped as tectonic
klippe perched anomalously above other members of theYuli Belt (Lin 1999b; Yen 1983; Yang and Wang 1985).
Finally, a gently dipping detachment with a normal senseof shear would accentuate exhumation cooling.
6.5 Implications for topographic growth

Considering the possible importance of both tectonic
and surface processes in exhumation, we evaluate thepossibility that these processes outpaced rock uplift andlimited the topographic growth of the orogen, at least to
the end of the youngest, penetrative deformation in the
Tailuko and Yuli Belts, ~0.7 Ma. We evaluate this possibility by showing a correlation between the age of thelate-stage deformation in the metamorphic rocks andthe results of a recent study documenting the paleotopography of the Backbone Range (Fig. 30) (Chang et al.
2023). The paleotopography analysis used the H isotope
record of n-alkanes with 29 carbon atoms (δ2HnC29) preserved in leaf waxes collected from exposed sedimentaryrocks in the Coastal Range. The sediments form part ofan uplifted forearc basin that accumulated sediments
derived from both the magmatic arc to the east and theemerging orogenic belt to the west from ~4.0 to 0.7 Ma.
A younger terrace deposit < 0.5 Ma was also sampled.
The δ2HnC29 values are consistently high from 3.0 to
1.25 Ma with a mean of -160 (0/00) and decrease slightly

Fig. 30 Modeled elevation for the source of organic materials
versus age (from Chang et al. 2023). The age range of late-stage
deformation in the eastern Backbone Range is based on isotopic ages
and bending of the northern Backbone Range orocline (this study)

Page 45 of 50

to a mean of ~− 180 (0/00) ~1.0 Ma. A final drop to − 200(0/00) at 0.3 Ma (Fig. 30). Chang et al. (2023) modeled the
H data using a modern Taiwan reference and proposedthat mean source elevation grew from a few hundred
meters 3.0–1.25 Ma, to 1.5 km ~1.0 Ma and reached
its current mean elevation of just over 2.5 km < 0.5 Ma(Fig. 30). These data strongly argue that topographic
growth was limited despite clear evidence of significant
erosion (e.g., high deposition rates) during this time. One
interpretation, therefore, is that tectonic thinning of theeastern Backbone Range played an important role in limiting topographic growth.
A relatively low elevation for the Backbone Range
until very recently is also consistent with preserving
areas of anomalously low relief in the Backbone Range,some perched at high elevations today. The areas ofanomalous topographic relief were first described by
Japanese researchers a century ago (Lin 1957 and references therein). They ascribed the surfaces to the decay
and rejuvenation of tectonic activity, generally following
ideas proposed by Davis (1899). Lin (1957) resurrected
this hypothesis, and Ouimet et al. (2013), based on 10Be
isotope data from the surfaces and the steep scarps often
surrounding the surfaces, documented extremely low
erosion rates on the surfaces and extremely high rates on
the eroding flanks. Yang et al. (2015) proposed that the
low-relief surfaces could form during stream piracy andheadwater migrations, although Whipple et al. (2017)questioned the significance of these processes. Morerecently, Hsieh et al. (2018) presented one of the mostdetailed studies of one of the largest surfaces in Taiwan,
the Alishan surface, and showed that the surface formed> 50 ky ago (the maximum age of 14C), probably during
the last glacial period. They propose that during relatively dry glacial periods, river downcutting decreases
and hillslopes degrade, potentially forming areas of relatively low relief. Similar processes may have been operating throughout Taiwan, including the eastern Backbone
Range, where tectonic exhumation may have played arole in limiting surface uplift. However, the low-relief
surfaces in the Backbone Range remain undated, andadditional research is needed to verify their origin (Additional file 16).

7 Conclusions
Integration of previous studies of the metamorphic corein the Taiwan orogen with new structural, geochronologic, and thermometric data suggests a new interpretation for the arc-continent collision. We propose a
two-stage process for exhuming the high-pressure metamorphic rocks of the core of the orogen. In the initialstage, exhumation is driven primarily by pressure gradients in a subduction channel, possibly enhanced byByrne et al. Progress in Earth and Planetary Science(2024) 11:23

partial subduction of continental crust. During this stage,
exhuming rocks are tectonically extruded up the subduction channel with vertical and horizontal components.
The horizontal component drives the second stage. During this stage, a previously recognized, margin-parallel,
left-lateral strike-slip system is offset, forming a left-stepping extensional bend as oblique convergence continued.
The extruding high-pressure rocks filled this extensionalzone as the hanging wall was structurally thinned.
The primary evidence for the new interpretation is theregional distribution of a late-stage, gently dipping penetrative foliation and a suite of sub-vertical veins. Isotopic
dating and cross-cutting relations indicate this exhumation stage extended from ca. 2.4 to < 1.0 Ma when the
Backbone Range orocline started to form. Combined tectonic and erosional processes may have limited the topographic growth of the orogen from ~3.0 to < 1.0 Ma. We
propose that the late-stage, gently dipping foliation andsteeply dipping veins formed in the footwall of a regionalscale, extensional shear zone rooted to the east beneaththe forearc (Fig. 29B). Normal slip on the detachment
accounts for CW rotation (north view) of older foliations
that represent the eastern half of an apparent cleavage fan
in the Backbone Range (red dashed line in Fig. 29B, crosssection) and the local occurrence of top-to-the-east shear
previously interpreted as evidence for back thrusting.
The proposed hypothesis could be further developed
and tested with additional structural and geochronological data. In particular, the southern Backbone Range,
where the margin-parallel strike-fault appears to be offset, is ripe for more detailed, integrated studies. Perhaps
one of the most challenging areas of future study is identifying and mapping remnants of the proposed extension
shear zone that accommodated the westward extrusionof the orogenic wedge. Future studies incorporating lowtemperature geochronometers and surface processes
could also better define the link between the inactive
structures described here and active structures thatdefine the tectonic regimes of modern Taiwan. Finally,
we assumed highly oblique plate convergence throughout the different deformation regimes because it is consistent with strike-slip deformation during D2, northeast
extension during D4, and recent reconstructions based
on tomographic data (Wu et al. 2016; Sibuet et al. 2021).
Presumably, highly oblique convergence transitioned
to the modern, more orthogonal convergence before
the northern Backbone Range orocline was initiated ~1
Mya. Linking plate motions to past and modern tectonic
regimes is also an area ripe for more detailed studies.
Abbreviations
AFT	Fission track dating of apatite grains
AHe	U-Th/He dates of apatite grains
C	Centigrade

Page 46 of 50

CCW​	Counterclockwise
CW	Clockwise
CF	Chaochou fault
CL	Cathodoluminescence
CR	Coastal range
CXIH	Central Cross-Island Highway
Dx	Deformation event with x = 1, 2, 3, or 4, reflecting the relative
age of different deformation events that include various tec‑
tonic structures, except S0, which represents bedding
EUR	Eurasian Plate
FIM	Fluid inclusion microthermometry
L/V	Liquid/vapor
Ma	Millions of years before present
Myr	Millions of years
PCV	Plate convergence vectorPSP	Philippine Sea Plate
RA	Ryukyu arc
RSCM	Ramen spectrometry of carbonaceous material
SEM-BSE	Scanning electron microscope and backscattered electrons
SEM-CL	Scanning electron microscope with cathodoluminescence
SEM–EDS	Scanning electron microscope and energy-dispersive X-ray
spectroscopy
Sx (italics S)	
Stretch direction with S = 1 + e and e = (final length-initiallength)/initial length and S1 > S2 > S3
Sx	Foliations with x = 0, 1, 2, 3, or 4, reflecting the relative age of dif‑ferent foliations
SXIH	South Cross-Island Highway
TSZ	Tailuko Shear Zone
X, Y, and Z	Principal strain axes
XPL	Cross-polarization
ZrnFT	Fission track dating of zircon grains
ZrnHe	U-Th/He dating of zircon grains

Supplementary Information
The online version contains supplementary material available at https://​doi.​
org/​10.​1186/​s40645-​024-​00627-w.
Additional file 1. Appendix I: Methods for collecting samples and struc‑tural and petrographic analyses.
Additional file 2. Appendix II: Methods for fluid inclusion microther‑mometry (FIM).
Additional file 3. Appendix II: FIM Table 1a, fluid inclusion petrographic
and microthermometric observations.
Additional file 4. Appendix II: FIM Table 1b, fluid inclusion petrographic
and microthermometric observations.
Additional file 5. Appendix II: FIM Table 1c, fluid inclusion petrographic
and microthermometric observations.
Additional file 6. Appendix II: FIM Table 2a, fluid inclusion microther‑mometry results as calculated in fluids.
Additional file 7. Appendix II: FIM Table 2b, fluid inclusion microther‑mometry results as calculated in fluids.
Additional file 8. Appendix II: FIM Table 2c, fluid inclusion microther‑mometry results as calculated in fluids.
Additional file 9. Appendix II: Description of maps used to approximate
the shape of each sample chip used for FIM.
Additional file 10. Appendix II: Map 1, sample T19H-Q, northerndomain.
Additional file 11. Appendix II: Map 2, sample T19H-ADL, northerndomain.
Additional file 12. Appendix II: Map 3, sample S1-A (quartz), central
domain.
Additional file 13. Appendix II: Map 4, sample S1-B (quartz), central
domain.

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Additional file 14. Appendix II: Map 5, sample Xin05 (quartz), southerndomain.
Additional file 15. Appendix III: 40Ar/39Ar step-heating and inverse
isochrons for the chlorite-coated adularia sample T19HC-03.
Additional file 16. Appendix III: 40Ar/39Ar data tables for adularia
sample 19THC-01.
Acknowledgements
The authors are extremely grateful for the support and encouragement fromour friends and colleagues in Taiwan. This has been a multi-year project withmany logistical and scientific challenges. We are especially grateful to Dr.
Wen-Yen Chang, Dean of the School of Environment at National Dong Hwa
University, who helped secure a very comfortable office and apartment for TB
in 2021. A long-weekend car rental was also appreciated. Graduate students at
NDHU were especially inviting and patient during research meetings in theirresearch lab. Chih-Ying Yeh, in particular, was a great help as she transitioned
from graduate student to lab technician. Ph.D. students Wen-Han Lo andDominikus Deka Dewangga at NDHU kindly shared their research ideas andobservations, and Wen-Han led a critical expedition to the deepest parts of
the high-grade rocks along the Chingsuei River. Dr. Sun-Lin Chung, Directorof the Institute of Earth Sciences, provided much-needed logistical supportand encouragement to work on “S3” in the Yuli Belt. Wei-Hao Hsu and ChiouLien Chang “Zinc,” graduate students at National Taiwan University, helpedwith collecting samples at Kukutsu Cliffs. Chi-Hsiu Pang and Kai-Hsiang Yang,
graduate students at the Institute of Earth Sciences helped in the field on
several occasions and provided stimulating discussions on the geology of
Taiwan, and the Yuli Belt in particular. Lindsey Adam Cromwell, Lauren Donati,
Ross Bolesta, Susie Adams, Morgan Spatz, and Ben Falvo, all enthusiastic
undergraduate students at IUP, assisted in the field for at least two fieldseasons while supported by NSF award EAR1650157 to JCL. Many additional
undergraduate students from universities and academic institutions in Taiwan
also participated in the field expeditions and always provided comfortableand entertaining field excursions. Professional field guides secured our safety
on many occasions and made delicious meals, often on the bank of the riverin the middle of nowhere. Rick Allmendinger’s Stereonet program was an
essential tool throughout the study and is greatly appreciated. Blake Restelli,
RJ Lee Group, and Emily Peterman, Bowdoin College, provided professional
guidance in preparing sections for SEM and CL analysis. Christopher McDon‑
ald, Arizona State University, provided critical assistance in interpreting the Ar/
Ar geochronology.
Author contributions
TB, JCL, and J-CL conceived and designed various aspects of the research. MCcollected field and thin-section data and analyzed the fluid inclusions. G-RH
collected field data, led field expeditions, and participated in data interpreta‑
tions. Y-HL and E-CY collected field data and provided critical insights into thestructural history of the Yuli and Tailuko Belts. C-H Tsai provided expertise inmetamorphic petrology and assistance and guidance in the SEM lab at NDHU.
ME assisted in fluid inclusion analysis and interpretation. LW collected andanalyzed the 40Aar/39Ar data. TB led the writing with contributions from MC,
JW, J-CL, ME, C-HT, and LW.
Funding
The research was supported by Grants from the Ministry of Science and Tech‑
nology (MOST) of Taiwan (MOST 107-2116-M-001-026-MY3; MOST 110-2116M-001-004 to Jian-Cheng Lee and MOST 109-2811-M-259-513 to Chin-Ho Tsai)and the US National Science Foundation (EAR-1220453 to Byrne; EAR-1650157
to Lewis).
Availability of data and materials
The article and the appendices include the data supporting the conclusions
presented here.

Declarations
Competing interests
The authors declare no competing interests.

Page 47 of 50

Received: 4 September 2023 Accepted: 6 April 2024

References
Avé Lallemant HG, Guth LR (1990) Role of extensional tectonics in exhuma‑tion of eclogites and blueschists in an oblique subduction setting:northwestern Venezuela. Geology 18:950–953. https://​doi.​org/​10.​1130/​0091-​7613(1990)​018%​3c0950:​ROETIE%​3e2.3.​CO;2
Baziotis I, Tsai CH, Ernst WG, Jahn BM, Iizuka Y (2016) New P-T constraints on
the Tamayen glaucophane-bearing rocks, eastern Taiwan: Perple_X
modelling results and geodynamic implications. J Metamorph Geol
35:35–54
Beyssac O, Simoes M, Avouac J-P, Farley K, Chen Y-G, Chan Y-C, Goffe B (2007)
Late Cenozoic metamorphic evolution and exhumation of Taiwan.
Tectonics 26:TC6001. https://​doi.​org/​10.​1029/​2006T​C0020​64
Bons PD (2000) The formation of veins and their microstructures. In: Stress,
strain and structure: a volume in honor of W. D. Means, vol 2. Australia.
https://​doi.​org/​10.​3809/​jvirt​ex.​2000.​00007
Chan Y-C, Okamoto K, Yui T-F, Iizuka Y, Chu H-T (2005) Fossil fluid reservoir
beneath a duplex fault structure within the Central Range of Taiwan:
implications for fluid leakage and lubrication during earthquake ruptur‑
ing process. Terra Nova 17(6):493–499. https://​doi.​org/​10.​1111/j.​1365-​
3121.​2005.​00636.x
Chang C-P, Angelier J, Lu C-Y (2007) Polyphase deformation in a newly
emerged accretionary prism: folding faulting and rotation in the south‑
ern Taiwan mountain range. Tectonophysics. https://​doi.​org/​10.​1016/j.​
tecto.​2007.​11.​002
Chang CP, Angelier J, Huang CY (2009) Evolution of subduction indicated bymelanges in Taiwan. In: Lallemand S, Funiciello F (eds) Subduction zone
geodynamics. Springer, Berlin, pp 207–225. https://​doi.​org/​10.​1007/​
978-3-​540-​87974-9
Chang Q, Hren M, Lai LS-H, Dorsey R, Byrne T (2023) Rapid topographic growthof the Taiwan orogen since 1.5 Ma. Sci Adv. https://​doi.​org/​10.​1126/​
sciadv.​ade64​15
Chemenda AI, Mattauer M, Malavieille J, Bokun A (1995) A mechanism forsyn-collisional rock exhumation and associated normal faulting: Resultsfrom physical modelling. Earth Planet Sci Lett 132:225–232
Chen C-H (2000) Geologic Map of Taiwan. Central Geological Survey, Ministryof Economic Affairs
Chen WS, Chen YC, Yu NT, Yen IC, Yang BCC, Shih TS (2005) Stratigraphic andtectonic geomorphic studies along the Chao Chou fault, Southern
Taiwan-Reevaluating fault characteristics and location. Spec Publ Cent
Geol Surv 16:75–90
Chen C-T, Wu C-Y, Lo C-H, Chu H-T, Yui T-F (2017a) Dating palaeo-seismic fault‑
ing in the Taiwan Mountain Belt. Terra Nova 30:146–151. https://​doi.​
org/​10.​1111/​ter.​12319
Chen WS, Chung S-L, Chou H-Y, Zugererbai Z, Shao W-Y, Lee Y-H (2017b) A
reinterpretation of the metamorphic Yuli belt: evidence for a middlelate Miocene accretionary prism in eastern Taiwan. Tectonics. https://​doi.​org/​10.​1002/​2016T​C0043​83
Chen W-S, Yeh J-J, Syu S-J (2019) Late Cenozoic exhumation and erosion of
the Taiwan orogenic belt: new insights from petrographic analysis
of foreland basin sediments and thermochronological dating on themetamorphic orogenic wedge. Tectonophysics 750:56–69. https://​doi.​
org/​10.​1016/j.​tecto.​2018.​09.​003
Chen C-T, Lo C-H, Wang P-L, Lin L-H (2022) Extensional mountain building
along convergent plate boundary: insights from the active Taiwan
mountain belt. Geology 50:1245–1249. https://​doi.​org/​10.​1130/​
G50311.1
Clark MB, Fisher DM, Lu C-Y (1992) Strain variations in the Eocene and older
rocks exposed along the central and southern cross-island highways.
Taiwan Acta Geol Taiwan 30:1–10
Conand C, Mouthereau F, Ganne J, Lin AT-S, Lafhid A, Daudet M, Mesalles L,
Giletycz S, Bonzani M (2020) Strain partitioning and exhumation in
Taiwan oblique collision: role of rift architecture and plate kinematics.
Tectonics. https://​doi.​org/​10.​1029/​2019T​C0057​98
Craw D, Upton P, Yu B-S, Horton T, Chen Y-G (2009) Young orogenic gold
mineralisation in active collisional mountains, Taiwan. Miner Depos
45:631–646

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Crespi J, Chan Y-C, Swaim M (1996) Synorogenic extension and exhumation of
the Taiwan hinterland. Geology 24(3):247–250
Dadson SJ, Hovius N, Chen H, Dade WB, Hsieh M-L, Willett SD, Hu J-C, Horng
M-J, Chen M-C, Stark CP, Lague D, Lin J-C (2003) Links between ero‑sion, runoff variability and seismicity in the Taiwan orogen. Nature
426:648–651
Davis WM (1899) The geographical cycle. Geogr J 14:481–504
Derrieux F, Siame L, Bourlés D, Chen R-F, Braucher R, Léanni L, Lee J-C, Chu H-T,
Byrne T (2014) How fast is the denudation of the Taiwan Mountain belt?
J East Asian Sci. https://​doi.​org/​10.​1016/j.​jseaes.​2014.​03.​012
Dewangga DD, Tsai C-H, lizuka Y, Kouketsu Y, Sakaguchi IT, Chung S-L, Chen
H-F (2021) Origins and metamorphism of serpentinites in the Yuli belt,
eastern Taiwan. In: Geologica Society of Taiwan, Taipei
Ernst WG (1977) Olistostromes and included ophiolitic debris from the Coastal
Range of eastern Taiwan. Mem Geol Soc China 2:97–114
Faure M, Lu C-Y, Chu H-T (1991) Ductile deformation and Miocene nappestacking in Taiwan related to motion of the Philippine Sea Plate.
Tectonophysics 198:95–105
Feehan JG, Brandon MT (1999) Contribution of ductile flow to exhumation of
low-temperature, high-pressure metamorphic rocks: San Juan, Cascade
nappes, NW Washington State. J Geophys Res 104:10883–10902
Fitch TJ (1972) Plate convergence, transcurrent faults, and internal deforma‑tion adjacent to Southeast Asia and the Western Pacific. J Geophys Res
77(23):4432–4460
Fossen H (2010) Structural geology. Cambridge University Press, Cambridge
Fossen H, Cavalcante GCG (2017) Shear zones—a review. Earth Sci Rev
171:434–455
Fuller C, Willett SD, Fisher DM, Lu C-Y (2006) A thermomechanical wedge
model of Taiwan constrained by fission-track thermochronometry.
Tectonophysics 425:1–24
Grujic D, Casey M, Davidson C, Hollister LS, Kündig R (1996) Ductile extrusionof the Higher Himalayan Crystalline in Bhutan: evidence from quartz
microfabrics. Tectonophysics 260:21–43
Ho CS (1986) A synthesis of the geologic evolution of Taiwan. Tectonophysics
125:1–16
Ho CS (1988) An introduction to the geology of Taiwan–explanatory text ofthe geologic map of Taiwan, 2nd edn. Ministry of Economic Affairs,
TaipeiHo G-R, Lo W (2015) Structural analysis in Shoufengsi area of Tananao complex,
Eastern Taiwan. Terr Atmos Ocean Sci 26(5):557–569
Ho G-R, Byrne TB, Lee J-C, Mesalles L, Lin C-W, Lo W, Chang C-P (2022) A
new interpretation of the metamorphic core in the Taiwan orogen:a regional-scale, left-lateral shear zone that accommodated highly
oblique plate convergence in the Plio-Pleistocene. Tectonophysics
833:229332. https://​doi.​org/​10.​1016/j.​tecto.​2022.​229332
Hsieh M-L, Hogg A, Kang S-C, Chou C-Y (2018) The preservation of lastglacial (N50 to 40 ka) colluvium on low-relief surfaces in Alishan, an
actively uplifting mountain in southwestern Taiwan. Geomorphology
322:159–174. https://​doi.​org/​10.​1016/j.​geomo​rph.​2018.​08.​041
Hsieh AI, Dashtgard SE, Wang P-L, Horng C-S, Su C-C, Lin AT, Vaucher R,
Löwemark L (2023) Multi-proxy evidence for rapidly shifting sedimentsources to the Taiwan Western Foreland Basin at the Miocene-Pliocenetransition. Basin Res 35(3):932–948. https://​doi.​org/​10.​1111/​bre.​12741
Hsu W-H, Byrne TB, Ouimet W, Lee Y-H, Chen Y-G, van Soest M, Hodges K(2016) Pleistocene onset of rapid, punctuated exhumation in the east‑
ern Central Range of the Taiwan orogenic belt. Geology 44(9):719–722.
https://​doi.​org/​10.​1130/​G37914.1
Hu J-C, Yu S-B, Angelier J (2002) Transition tectonics of northern Taiwan
induced by convergence and trench retreat. Geol Soc Am Spec Pap.
https://​doi.​org/​10.​1130/0-​8137-​2358-2.​147
Huang CK (1953) Adularia from Kukutzu, Taiwan. Acta Geol Taiwanica 5:27–34
Huang C, Byrne T (2014) Tectonic evolution of an active tectonostratigraphic
boundary in accretionary wedge: an example from the TulungwanChaochou Fault system, southern Taiwan. J Struct Geol. https://​doi.​org/​
10.​1016/j.​jsg.​2014.​06.​007
Huang C-Y, Yuan P, Tsao S-J (2006) Temporal and spatial records of active
arc-continent collision in Taiwan: a synthesis. Geol Soc Am Bull
118(3/4):274–288. https://​doi.​org/​10.​1130/​B25527.1
Huang H-H, Wu YM, Song X, Chang C-H, Lee S-J, Chang T-M, Hsieh H-H(2014) Joint Vp and Vs tomography of Taiwan: implications forPage 48 of 50

subduction-collision orogeny. Earth Planet Sci Lett 392:177–191.
https://​doi.​org/​10.​1016/j.​epsl.​2014.​02.​026
Huang P, Wei C, Dong J, Zhang J (2023) Two-stage exhumation of high–P
rocks from the Yuli Belt, Eastern Taiwan: insights from the metamorphic
evolution in subduction channels. Lithos 440–441:107056. https://​doi.​
org/​10.​1016/j.​lithos.​2023.​107056
Jahn BM, Liou JG, Nagasawa H (1981) High-pressure metamorphic rocks of
Taiwan: REE geochemistry, Rb–Sr ages and tectonic implications. Mem
Geol Soc China 4:497–520
Jiao ZH (1991) The geologic structure and evolution of the hoping area in
Eastern Taiwan. National Taiwan University, Taipei
Keyser W, Tsai CH, Iizuka Y, Oberhänsli R, Ernst WG (2016) High-pressure meta‑morphism in the Chinshuichi area, Yuli belt, eastern Taiwan. Tectono‑
physics 692:191–202
Kirstein L, Fellin M, Willett S, Carter A, Chen Y-G, Garver J, Lee D (2009) Pliocene
onset of rapid exhumation in Taiwan during arc-continent collision:
new insights from detrital thermochronometry. Basin Res 22:270–285.
https://​doi.​org/​10.​1111/j.​1365-​2117.​2009.​00426.x
Korren C, Ferre E, Yeh E-C, Chou Y-M, Chu H-T (2017) Seismic rupture param‑
eters deduced from a pliocene-pleistocene fault pseudotachylyte in
Taiwan. In: Thomas MY, Mitchell TM, Bhat HS (eds) Fault zone dynamic
processes: evolution of fault properties during seismic rupture, volGeophysical Monograph 227. Wiley, New York, pp 21–35
Kouketsu Y, Tsai CH, Enami M (2019) Discovery of unusual metamorphic
temperatures in the Yuli belt, eastern Taiwan: new interpretation ofdata by Raman carbonaceous material geothermometry. Geology
47(6):522–526. https://​doi.​org/​10.​1130/​G45934.1
Kuiper KF, Deino A, Hilgen FJ, Krijgsman W, Renne PR, Wijbrans AJ (2008) Syn‑chronizing rock clocks of Earth history. Science 320(5875):500–504
Lallemand S, Liu C-S, Dominguez S, Schnürle P, Malavieille J, Crew AS (1999)
Trench-parallel stretching and folding of forearc basins and lateralmigration of the accretionary wedge in the southern Ryukyus: acase of strain partition caused by oblique convergence. Tectonics
18(2):231–247
Lallemand S, Theunissen T, Schnürle P, Lee C-S, Liu C-S, Font Y (2013) Indenta‑tion of the Philippine Sea plate by the Eurasia plate in Taiwan: details
from recent marine seismological experiments. Tectonophysics. https://​doi.​org/​10.​1016/j.​tecto.​2013.​03.​020
Lan C-Y, Liou JG (1981) Occurrence, petrology and tectonics of serpentinites
and their associated rodingites in the Central Range, Taiwan. Mem Geol
Soc China 4:343–389
Lan C-Y, Liou JG (1984) Mineral chemistry of metamorphosed ocean rocks
in the Yuli belt of the Tananao schist. Taiwan Mem Geol Soc China
6:153–179
Lee Y-H (1995) A study of the stress field system evolution of the Taiwan
mountain belt, Central Cross-Island Highway and surrounding areas.
Trans Cent Geol Surv, MOEA, Taiwan 10:51–89
Lee C-T, Wang Y (1988) Quaternary stress changes in northern Taiwan and theirtectonic implication. Proc Geol Soc China 31:154–168
Lee Y-H, Yang CN (1994) Structural evolution in the Tayulin area of the central
range. Bull Cent Geol Surv 9:77–106
Lee JY, Marti K, Severinghaus JP, Kawamura K, Yoo HS, Lee JB, Kim JS (2006) A
redetermination of the isotopic abundances of atmospheric Ar. Geo‑chim Cosomochim Acta 70(17):4507–4512
Lee Y-H, Byrne T, Lo W, Wang S-J, Tsao S-J, Chen C-H (2022a) Out of sequencefaulting in backbone range and its implication to exhumation processof the Taiwan Orogen. In: Japan Geosciences Union/American Geo‑physical Union, Chiba, Japan (hybrid). vol SIT20-10. Japan Geosciences
Union
Lee Y-H, Byrne TB, Lo W, Wang S-J, Tsao S-J, Chen C-H, Yu H-C, Tan X, vanSoest M, Hodges K, Mesalles L, Robinson H, Fosdick JC (2022b) Outof sequence faulting in the backbone range, Taiwan: implications for
thickening and exhumation processes. Earth Planet Sci Lett 594:117711.
https://​doi.​org/​10.​1016/j.​epsl.​2022.​117711
Lewis JC, O’Hara DJ, Rau R-J (2015) Seismogenic strain across the transition
from fore-arc slivering to collision in southern Taiwan. J Geophys Res:
Solid Earth 120(6):4539–4555. https://​doi.​org/​10.​1002/​2015J​B0119​06
Lin C-C (1957) Topography of Taiwan. Publication of the Taiwan Provincial
Documentary Committee, Taipei (in Chinese)Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Lin M-L (1999a) Existence of the Chouochi transverse fault reconfirmed insoutheastern Taiwan and its tectonic implication. J Geol Soc China
42(1):69–84
Lin M-L (1999b) Litho-Stratigraphy and structural geology of the Wanjungarea, eastern Taiwan and their tectonic implications. J Geol Soc China
42(2):247–267
Lin CW, Chen W-S (2016) Geololgic map of Taiwan. Geological Society of
Taiwan, Taiiwan
Lin M-L, Yang CN, Wang Y (1984) Petrotectonic study on the Yuli Belt of the
Tananao Schist in the Chinshuichi area, Eastern Taiwan. Acta Geol
Taiwan 22:151–188
Liou JG (1981) Petrology of metamorphosed oceanic rocks in the Central
Range of Taiwan. Mem Geol Soc China 4:281–342
Liou JG, Lan CY, Suppe J, G. EW, (1977) The East Taiwan Ophiolite, its occur‑
rence, petrology, metamorphism and tectonic setting. Min Res Serv
Organ Spec Rep 1:212p
Liu T-K (1982) Tectonic implication of fission track ages from the Central Range,
Taiwan. Proc Geol Soc China 25:22–37
Liu T-K, Hsieh S, Chen Y-G, Chen W-S (2001) Thermo-kinematic evolution of the
Taiwan oblique-collision mountain belt as revealed by zircon fission
track dating. Earth Planet Sci Lett 186:45–56
Liu YC, Chen PT, Huang C, Yen IC, Lu ST (2009) Resent survey results of Chao‑chou fault. Annu Rep Cent Geol Surv 2008:35–36
Lo C-H, Onstott TC (1995) Rejuvenation of K-Ar systems for minerals in the
Taiwan Mountain Belt. Earth Planet Sci Lett 131:71–98
Lo C-H, Yui T-F (1996) 40Ar/39Ar dating of high-pressure rocks in the Tananao
basement complex, Taiwan. J Geol Soc China 39:13–30
Lo Y-C, Chen C-T, Lo C-H, Chung S-L (2020) Ages of ophiolitic rocks along plate
suture in Taiwan orogen: fate of the South China Sea from subduction
to collision. Terr Atmos Ocean Sci 31:383–402. https://​doi.​org/​10.​3319/​
TAO.​2019.​06.​19.​01
Lu C-Y, Lee J-C (2001) Structural evolution of the southeastern Central Range,
Taiwan. West Pac Earth Sci 1(2):213–226
Lu C-Y, Wang Lee C (1986) The sheath folds in the Tananao group between
Tienshiang and Tailuko, east-west cross-island highway, Taiwan.
Tectonophysics 125(1):125–131. https://​doi.​org/​10.​1016/​0040-​1951(86)​90010-7
Ludwig KR (2003) User’s manual for isoplot 3.00: a geochronological toolkit for
microsoft excel
Malavieille J (2010) Impact of erosion, sedimentation, and structural heritageon the structure and kinematics of orogenic wedges: analog models
and case studies. GSA Today. https://​doi.​org/​10.​1130/​GSATG​48A.1
Malavieille J, Trullenque G (2009) Consequences of continental subduction
on forearc basin and accretionary wedge deformation in SE Taiwan:
insights from analogue modeling. Tectonophysics 466(3–4):377–394.
https://​doi.​org/​10.​1016/j.​tecto.​2007.​11.​016
Mann P, Gordon MB (1996) Tectonic uplift of blueschist belts along transpres‑sional strike-slip faults. In: Bebout GE, Scholl DW, Kirby SH, Platt JP (eds)
Subduction top to bottom, vol 96. Am. Geo. U. Geophys. Monograph,
96 edn. AGU, pp 143–154
McCaffrey R (1992) Oblique plate convergence, slip vectors, and forearc defor‑
mation. J Geophys Res 97(B6):8905–8915
Mccaffrey R (2002) Crustal block rotations and plate coupling. In: Plate bound‑ary zones, pp 101–122. https://​doi.​org/​10.​1029/​GD030​p0101
Mesalles L (2014) Mountain building at a subduction-collision transition zone,
Taiwan: insights from morphostructural analysis and thermochro‑nological dating. Université Pierre et Marie Curie - Paris VI, Creative
Commons, Paris
Mesalles L, Mouthereau F, Bernet M, Chang C-P, Tien-Shun Lin A, Fillon C, Sen‑gelen X (2014) From submarine continental accretion to arc-continent
orogenic evolution: the thermal record in southern Taiwan. Geology
42(10):907–910. https://​doi.​org/​10.​1130/​g35854.1
Mesalles L, Lee Y-H, Ma T-C, Tsai W-L, Tan X-B, Lee H-Y (2020) A Late-Miocene
Yuli belt? New constraints on the eastern Central Range depositional
ages. Terr Atmos Ocean Sci 31(4):403–414. https://​doi.​org/​10.​3319/​TAO.​
2019.​06.​24.​01
Min K, Mundil R, Renne PR, Ludwig KR (2000) A test for systematic errors in
40Ar/39Ar geochronology through comparison with U/Pb analysis of a
1.1-Ga rhyolite. Geochim Cosomochim Acta 64(1):73–98

Page 49 of 50

Mondro C, Fisher D, Yeh E-C (2017) Strain histories from the eastern Central
Range of Taiwan: a record of advection through a collisional orogen.
Tectonophysics. https://​doi.​org/​10.​1016/j.​tecto.​2017.​03.​007
Osozawa S, Shinjo R, Armid A, Watanabe Y, Horiguchi T, Wakabayashi J (2012)
Palaeogeographic reconstruction of the 1.55 Ma synchronous isolationof the Ryukyu Islands, Japan, and Taiwan and inflow of the Kuroshio
warm current. Int Geol Rev 54(12):1369–1388. https://​doi.​org/​10.​1080/​
00206​814.​2011.​639954
Ouimet W, Byrne T, Siame L, Bierman P, Rood DH (2013) Slow erosion rates,
increasing relief and transient landscape evolution within the Central
Range of Taiwan. In: Fall Meeting. American Geophysical Union, San
Francisco
Page BM, Suppe J (1981) The Pliocene Lichi melange in Taiwan: its plate tec‑
tonic and olistostromal origin. Am J Sci 281:193–227
Passchier C, Trouw RAJ (2005) Microtectonics, 2nd edn. Springer, Berlin.
https://​doi.​org/​10.​1007/3-​540-​29359-0
Pelletier B, Hu HN (1984) New structural data along two transects across thesouthern half of the Central Range of Taiwan. Mem Geol Soc China
6:1–19
Platt JP (1983) Progressive refolding in ductile shear zones. J Struct Geol
5(6):619–622
Platt JP (1993) Exhumation of high-pressure rocks: a review of concepts andprocesses. Terra Nova 5(2):119–133. https://​doi.​org/​10.​1111/j.​1365-​
3121.​1993.​tb002​37.x
Platt JP, Soto J-I, Whitehouse MJ, Hurford AJ, Kelley SP (1998) Thermal evolu‑tion, rate of exhumation, and tectonic significance of metamorphic
rocks from the floor of the Alboran extensional basin, western Mediter‑
ranean. Tectonics 17(5):671–689. https://​doi.​org/​10.​1029/​98TC0​2204
Pulver MH, Crespi JM, Byrne TB (2002) Lateral extrusion in a transpressional col‑lisional zone: an example from the pre-Tertiary metamorphic basementof Taiwan. In: Byrne TB, Liu C-S (eds) Geology and geophysics of anarc-continent collision, Taiwan. Geological Society of America, Special
Paper 358, pp 107–120
Ramsay JG (1980) The crack-seal mechanism of rock deformation. Nature
284:135–139
Rau R-J, Ching K-E, Hu J-C, Lee J-C (2008) Crustal deformation and block kin‑ematics in transition from collision to subduction: GPS measurements
in northern Taiwan, 1995–2005. J Geophys Res. https://​doi.​org/​10.​1029/​
2007J​B0054​14
Ring U, Brandon M (1999) Ductile deformation and mass loss in the Franciscan
Subduction Complex: implications for exhumation processes inaccretionary wedges. In: Ring U, Brandon M, Lister G, Willett S (eds)
Exhumation proceses: normal faulting, ductile flow and erosion, vol154. Geological Society of London, London, pp 55–86
Ring U, Kumerics C (2008) Vertical ductile thinning and its contribution to theexhumation of high-pressure rock: the Cycladic blueschist unit in the
Aegean. J Geol Soc Lond 165:1019–1030. https://​doi.​org/​10.​1144/​0016-​
76492​008-​010
Ring U, Brandon MT, Lister G, Willett S (1999) Exhumation processes: normal
faulting, ductile flow and erosion. Special Publication, vol 154. Geologi‑
cal Society of London, London
Rutte D, Becker TA, Deino AL, Reese SR, Renne PR, Schickler RA (2018) The new
CLOCIT irradiation facility for 40Ar/39Ar geochronology: characteri‑sation, comparison with CLICIT and implications for high-precisiongeochronology. Geostand Geoanal Res 42(3):301–307
Sandmann S, Nagel HJ, Froitzheim N, Ustaszewski K, Münker C (2015) Late
miocene to early pliocene blueschist from Taiwan and its exhumation
via forearc extraction. Terra Nova 27:285–291. https://​doi.​org/​10.​1111/​
ter.​12158
Secor DT (1965) Role of fluid pressure in jointing. Am J Sci 263:633–646
Shan Y, Nie G, Ni Y, Chang C-P (2014) Structural analysis of a newly emerged
accretionary prism along the Jinlun-Taimali coast, southeastern Taiwan:
from subduction to arc-continent collision. J Struct Geol 66:248–260.
https://​doi.​org/​10.​1016/j.​jsg.​2014.​06.​002
Shen T-T, Liu T-K, Huang S-Y, Hsieh P-S, Wu C-Y (2020) Post-collisional exhuma‑tion and geotherm pattern in northern Tananao Complex, northeastern
Taiwan. Terr Atmos Ocean Sci 31:369–381. https://​doi.​org/​10.​3319/​TAO.​
2019.​04.​06.​01
Shepherd TJ, Rankin AH, Alderton DHM (1985) A practical guide to fluid inclu‑
sion studies. Blackie, Barcelona

Byrne et al. Progress in Earth and Planetary Science(2024) 11:23

Shyu B, Sieh K (2005) Neotectonic architecture of Taiwan and its implications
for future large earthquakes. J Geophys Res. https://​doi.​org/​10.​1029/​
2004J​B0032​51
Sibuet J-C, Hsu S-K, Shyu C-T, Liu C-S (1995) Structural and kinematic evolu‑
tions of the okinawa trough backarc basin. In: Taylor B (ed) Backarc
basins: tectonics and magmatism. Springer, Boston, pp 343–379.
https://​doi.​org/​10.​1007/​978-1-​4615-​1843-3_9
Sibuet J-C, Deffontaines B, Hsu S-K, Thareau N, Le Formal J-P, Liu C-S (1998)
Okinawa trough backarc basin: early tectonic and magmatic evolution.
J Geophys Res: Solid Earth 103(B12):30245–30267. https://​doi.​org/​10.​1029/​98JB0​1823
Sibuet J-C, Zhao M, Wu J, Lee C-S (2021) Geodynamic and plate kinematic con‑text of South China Sea subduction during Okinawa trough opening
and Taiwan orogeny. Tectonophysics 817:229050. https://​doi.​org/​10.​1016/j.​tecto.​2021.​229050
Simoes M, Beyssac O, Chen Y-G (2012) Late Cenozoic metamorphism andmountain building 1 in Taiwan: a review. J Asian Earth Sci 46:92–119
Stanley RS, Hill AB, Chang HC, Hu HN (1981) A transect through the metamor‑
phic core of the Central Mountains, southern Taiwan. Mem Geol Soc
China 4:443–473
Stoenner RW, Schaeffer OA, Katcoff S (1965) Half-lives of argon-37, argon-39,
and argon-42. Science 148(3685):1325–1328
Suppe J (1981) Mechanics of mountain building and metamorphism in
Taiwan. Mem Geol Soc China 4:67–89
Suppe J (1984) Kinematics of arc-continent collision, flipping of subduction
and back-arc spreading near Taiwan. Mem Geol Soc China 6:21–33
Tan LP, Wang Lee C, Chen CC, Tien PL, Tsui PC, Yui TF (1978) A mineralogicalstudy of the Fengtien nephrite deposits of Hualien: Taiwan. Natl Sci
Counc Spec Publ 1:81
Teng LS (1990) Geotectonic evolution of late Cenozoic arc-continent collisionin Taiwan. Tectonophysics 183:57–76
Teng LS, Lee C-T, Peng C-H, Chen W-F, Chu C-J (2001) Origin and geologi‑
cal evolution of the Taipei Basin, northern Taiwan. West Pac Earth Sci
1(2):115–142
Tobisch OT, Paterson SR (1988) Analysis and interpretation of composite folia‑
tions in areas of progressive deformation. J Struct Geol 10(7):745–754.
https://​doi.​org/​10.​1016/​0191-​8141(88)​90081-8
Touret JLR (2001) Fluids in metamorphic rocks. Lithos 55:1–25
Tsai C-H, Iizuka Y, Ernst WG (2013) Diverse mineral compositions, textures,and metamorphic P-T conditions of the glaucophane-bearing rocks
in the Tamayen mélange, Yuli belt, eastern Taiwan. J Asian Earth Sci
63:218–233. https://​doi.​org/​10.​1016/j.​jseaes.​2012.​09.​019
Usami M, Matsumoto T (1940) Geologic map of Taiwan: Daibuzan Sheet.
Bureau of Produce Industry, Government Gen, Taiwan
Ustaszewski K, Wu Y-M, Suppe J, Huang H-H, Chang C-H, Carena S (2012)
Crust–mantle boundaries in the Taiwan–Luzon arc-continent collisionsystem determined from local earthquake tomography and 1D models:
Implications for the mode of subduction polarity reversal. Tectonophys‑
ics 578:31–49. https://​doi.​org/​10.​1016/j.​tecto.​2011.​12.​029
Van Avendonk HJA, McIntosh KD, Kuo-Chen H, Lavier LL, Okaya DA, Wu FT,Wang C-Y, Lee C-S, Liu C-S (2015) A lithospheric profile across northern
Taiwan: from arc-continent collision to extension. Geophys J Int
204(1):331–346. https://​doi.​org/​10.​1093/​gji/​ggv468
von Hagke C, Philippon M, Avouac JP, Gurnis M (2016) Origin and time evolu‑tion of subduction polarity reversal from plate kinematics of Southeast
Asia. Geology 44(8):659–662. https://​doi.​org/​10.​1130/​G37821.1
Wang PL (1998) K-feldspar40Ar/39Ar thermochronology and its applications.
National Taiwan University, Taipei
Wang PL, Lin LH, Lo CH (1998) 40Ar/39Ar dating of mylonitization in the
Tanano Schist, eastern Taiwan. J Geol Soc China 159–183:1
Wang Lee C (1979) Preliminary notes on the time relations of metamorphic
rocks, Loshao–Tailuko area, eastern Taiwan. Mem Geol Soc China
3:249–252
Wang Lee C, Lu C-Y (1984) Field trip guide to the eastern section of the Cross
Island Highway. In: Paper presented at the Field guidebook for the
Sino-French Colloquium on Geodynamics of the Eurasian-Philippine
Sea Plate Boundary, Taiwan
Warren CJ (2013) Exhumation of (ultra-)high-pressure terranes: concepts andmechanisms. Solid Earth 4:75–92. https://​doi.​org/​10.​5194/​se-4-​75-​2013
Warren CJ, Beaumont C, Jamieson RA (2008) Deep subduction and rapid exhu‑
mation: role of crustal strength and strain weakening in continental

Page 50 of 50

subduction and ultrahigh-pressure rock exhumation. Tectonics. https://​doi.​org/​10.​1029/​2008T​C0022​92
Wei K-Y, Chen Y-G, Liu T-K (1998) Sedimentary history of the Taipei Basinwith constraints from thermoluminescence dates. J Geol Soc China
41:109–125
Whipple KX, DiBiase RA, Ouimet WB, Forte AM (2017) Preservation or piracy:
diagnosing low-relief, high-elevation surface formation mechanisms.
Geology 45:91–94. https://​doi.​org/​10.​1130/​G38490.1
Willett S, Fisher DM, Fuller C, Yeh E-C, Lu C-Y (2003) Erosion rates and orgenicwedge kinematics in Taiwan inferred from fission-track thermochro‑nometry. Geology 31(11):945–948
Wu J, Suppe J, Lu R, Kanda R (2016) Philippine Sea and East Asian plate tecton‑
ics since 52 Ma constrained by new subducted slab reconstruction
methods. J Geophys Res. https://​doi.​org/​10.​1002/​2016J​B0129​23
Yang GS (1986) A geomorphological study of active faults in Taiwan–espe‑
cially on the relation between active faults and geomorphic surfaces.
Chinese Culture University, Taipei
Yang CN, Wang Y (1985) Petrotectonic study on the Yuli blet of the Tananao
schist in the Juisui area. East Taiwan Acta Geol Taiwanica 23:153–180
Yang R, Willett SD, Goren L (2015) In situ low-relief landscape formation as aresult of river network disruption. Nature 520:526–529
Yeh E-C, Wu Y-M, Shyu JBH, Chang C-H (2016) Tectonic implication of the 5th
March 2005, doublet earthquake in Ilan, Taiwan. Terr Atmos Ocean Sci
27:799–805
Yen TP (1963) The metamorphic belts within the Tananao schist terrain of
Taiwan. Proc Geol Soc China 6:72–74
Yen TP (1970) Structural analysis of the Slate Formation of Taiwan. Bull Geol
Surv Taiwan 21:1–54
Yen TP (1983) Late Cenozoic faulting in the Juisui area, eastern Taiwan. Bull Inst
Earth Sci Acad Sin 3:149–154
Yi D, Cheng ZY, Lin CW (2012) The geological sheet map “Guangfu,” scale
1:50,000, and its explanatory text. Central Geological Survey, MOEA,New Taipei CityYi D-C, Lin C-W, McAleer R, Kunk M, Wintsch R (2015) Exhumation of Metamor‑
phic rocks during the Taiwan Orogeny: A Study on the Soufeng Fault.
In: American Geophysical Union Fall Meeting. American GeophysicalUnion, San Francisco
Yui TF, Lu C-Y, Lo CH (1990) Tectonic evolution of the Tananao Schist Complexof Taiwan. In: Aubouin J, Bourgois J (eds) Tectonics of circum-Pacificcontinental margins. VSP, Utrecht, pp 193–209
Yui T-F, Huang FE, Xu J (1996) Raman spectrum of carbonaceous material a
possible metamorphic grade indicator for low-grade metamorphic
rocks. J Metamorph Geol 14:115–124
Yui TF, Okamoto K, Usuki T, Lan CY, Chu HT, Liou JG (2009) Late Triassic-Late
Cretaceous accretion/subduction in the Taiwan region along the east‑
ern margin of South China—evidence from zircon SHRIMP dating. Int
Geol Rev 51(4):304–328. https://​doi.​org/​10.​1080/​00206​81080​26363​69
Yui TF, Maki K, Lan CY, Hirata T, Chu HT, Kon Y, Yokoyama TD, Jahn BM, Ernst
WG (2012) Detrital zircons from the Tananao metamorphic complex of
Taiwan: implications for sediment provenance and Mesozoic tectonics.
Tectonophysics 541–543:31–42. https://​doi.​org/​10.​1016/j.​tecto.​2012.​
03.​013
Yui T-F, Usuki T, Chen C-Y, Ishida A, Sano Y, Suga K, Iizuka Y, Chen C-T (2014) Dat‑ing thin zircon rims by NanoSIMS: the Fengtien nephrite (Taiwan) is the
youngest jade on Earth. Int Geol Rev 56(16):1932–1944. https://​doi.​org/​
10.​1080/​00206​814.​2014.​972994
Zhang R, Li S, Suo Y, Liu J, Cao X, Zhou J, Jiang Z, Li X (2022) A forearc pull-apartbasin under oblique arc-continent collision: insights from the North
Luzon Trough. Tectonophysics 837:229461. https://​doi.​org/​10.​1016/j.​
tecto.​2022.​229461

Publisher’s NoteSpringer Nature remains neutral with regard to jurisdictional claims in pub‑lished maps and institutional affiliations.