Rapid conversion of an oceanic spreading center
to a subduction zone inferred from
Timothy E. Keenan
, John Encarnación
, Robert Buchwaldt
, Dan Fernandez
, James Mattinson
, and P. Benjamin Luetkemeyer
Department of Earth and Atmospheric Sciences, Saint Louis University, St. Louis, MO 63108;
Department of Earth and Environment, Boston University,
Boston, MA 02215;
Schlumberger-WesternGeco, Geosolutions-Interpretation, Houston, TX 77042;
Department of Earth Science, University of California,
Santa Barbara, CA 93106; and
Department of Geology and Environmental Earth Science, Miami University, Oxford, OH 45056
Edited by W. G. Ernst, Stanford University, Stanford, CA, and approved October 11, 2016 (received for review June 20, 2016)
Where and how subduction zones initiate is a fundamental tectonic
problem, yet there are few well-constrained geologic tests that
address the tectonic settings and dynamics of the process. Numerical
weakest parts of the plate tectonic system [Gurnis M, Hall C, Lavier L
(2004) Geochem Geophys Geosys 5:Q07001], but previous studies
have not favored them for subduction initiation because of the pos-
itive buoyancy of young lithosphere. Instead, other weak zones,
such as fracture zones, have been invoked. Because these models
differ in terms of the ages of crust that are juxtaposed at the site of
subduction initiation, they can be tested by dating the protoliths of
metamorphosed oceanic crust that is formed by underthrusting at
the beginning of subduction and comparing that age with the age
of the overlying lithosphere and the timing of subduction initiation
itself. In the western Philippines, we find that oceanic crust was less
than ∼1 My old when it was underthrust and metamorphosed at the
onset of subduction in Palawan, Philippines, implying forced sub-
duction initiation at a spreading center. This result shows that young
and positively buoyant, but weak, lithosphere was the preferred site
for subduction nucleation despite the proximity of other potential
weak zones with older, denser lithosphere and that plate motion
rapidly changed from divergence to convergence.
Subduction is the major driver of plate motion (1), and pro-
cesses at subduction zones are largely responsible for the
growth and evolution of continents through accretion, collision,
and magmatism. Despite its importance, the process of subduction
initiation is still debated partly because evidence from the early
stages of subduction is often obscured by later deformation and
magmatism. The lack of geologic constraints on how subduction
zones initiate remains a significant void in our understanding of
Subduction initiation has been addressed primarily through nu-
merical modeling (2–9). These studies have demonstrated the need
for a weak zone in the lithosphere to facilitate subduction. Based on
this, subduction has been proposed to initiate in a variety of settings
such as transform faults or fracture zones (10), passive continental
margins (6, 11), oceanic detachment faults (12), and oceanic
spreading centers (13). Spreading centers have been the least fa-
vored, however, because the lithosphere there is positively buoyant.
Two contrasting ideas regarding the dynamics (i.e., the forces) of
subduction initiation have been explored. In “spontaneous”sub-
duction initiation, a plate’s increasing density with age may even-
tually cause it to sink into the underlying asthenosphere (10, 14),
whereas in forced subduction initiation, external plate forces are
required to initiate subduction (3, 5, 15). Oceanic plates that are at
least ∼10 My old are negatively buoyant (16) and may undergo
either forced or spontaneous subduction initiation. Subduction
initiation within very young lithosphere near a spreading center,
however, can only be forced because the plate is still positively
buoyant. Interestingly, some numerical models predict that despite
the buoyancy of the plate and the ridge push force, the forces re-
quired to initiate underthrusting within the young, thin lithosphere
of a spreading center are lower than within older, thicker litho-
sphere, which requires increasingly larger forces to cause down-
bending of the stronger plate (1). Self-sustained subduction, driven
by a plate’s negative buoyancy, might eventually be achieved after
initiation at a spreading center, if forced convergence is sustained
until older, denser lithosphere finally enters the trench.
Well-constrained geologic tests of the aforementioned models
are necessary to carry the debate forward. Because transform faults,
fracture zones, and continental margins juxtapose lithosphere of
different ages, whereas plates of equal and approximately zero age
are adjacent at spreading centers, determining where subduction
has initiated may be possible by comparing the ages of the un-
derthrust and overriding lithosphere at the time of subduction
initiation, in relation to the time of subduction initiation itself.
The timing of subduction initiation in some paleo-subduction
zones may be determined by constraining the timing of high tem-
perature metamorphism—associated with the initiation of sub-
duction—of the uppermost portions (i.e., the crust) of the initial
subducted plate. This metamorphic material may be transferred to
(or “welded”) and preserved underneath the mantle peridotite
hanging wall of the nascent subduction zone forearc of the upper
plate as heat from the overlying mantle, and the resulting ductile
shearing, progressively propagates down into the cold underthrust
Subduction, the process by which tectonic plates sink into the
mantle, is a fundamental tectonic process on Earth, yet the
question of where and how new subduction zones form remains
a matter of debate. In this study, we find that a divergent plate
boundary, where two plates move apart, was forcefully and
rapidly turned into a convergent boundary where one plate
eventually began subducting. This finding is surprising because,
although the plate material at a divergent boundary is weak, it is
also buoyant and resists subduction. This study suggests that
buoyant, but weak, plate material at a divergent boundary can
be forced to converge until eventually older and denser plate
material enters the nascent subduction zone, which then
Author contributions: T.E.K. and J.E. designed research; T.E.K., J.E., R.B., D.F., J.M., C.R.,
and P.B.L. performed re search; R.B. contributed n ew reagents/analytic tool s; T.E.K., J.E.,
R.B., D.F., J.M., and C.R . analyzed data; J.E., D.F., C. R., and P.B.L. performed fieldwo rk;
and T.E.K., J.E., and R. B. wrote the paper.
The authors declare no conflict of interest.
This article is a PNAS Direct Submission.
To whom correspondence should be addressed. Email: email@example.com.
This article contains supporting information online at www.pnas.org/lookup/suppl/doi:10.
www.pnas.org/cgi/doi/10.1073/pnas.1609999113 PNAS Early Edition
AND PLANETARY SCIENCES
crust (17). These high-temperature “metamorphic soles”that form
during the initiation of subduction (18, 19) are found at the base of
ophiolites—sections of oceanic lithosphere now on land—that are
often thought to represent the trapped forearc (i.e., the upper plate)
of subduction zones (20, 21). The high-temperature (>700 °C)
amphibolite-granulite facies rocks preserved in metamorphic soles
(18, 22, 23) indicate that they formed during the very early stages
of subduction, before the development of depressed isotherms
that produce typical blueschist facies rocks during more mature
subduction (24). Also important, the high pressures (∼10 kbar)
(18, 22, 23) associated with the formation of some metamorphic
soles indicate that they formed at oceanic mantle depths during
subduction initiation rather than by some other process like oce-
anic core complex formation or intraoceanic thrusting unrelated
to subduction initiation.
Metamorphic soles commonly have metamorphic cooling ages
that are similar to the igneous crystallization ages of their overlying
ophiolites (19, 25, 26), implying that the overlying ophiolite was still
very young or formed during or shortly after subduction initiated.
However, because these cooling ages reflect the time of meta-
morphic cooling and not necessarily the time when the original ig-
neous crust formed, the similarity in ages between the metamorphic
sole and overlying ophiolite may be interpreted in several ways:
(i) subduction initiation between older lithospheric plates followed
by rapid slab rollback and seafloor spreading that generates the
ophiolite eventually preserved with the sole (10); (ii) subduction
initiation of variably older lithosphere beneath an already active
spreading center (27, 28); (iii) subduction initiation along weak
detachment faults at some distance from a spreading center (12), or
(iv) underthrusting of young lithosphere at a spreading center (13)
(Fig. 1). Without the age of the initially subducted plate, it may not
be possible to differentiate between these models.
The geochemical affinities of the ophiolite-sole pairs could vary
depending on where, and how, the subducting and overriding
plates were generated. Thus, the geochemistry of the sole and
overlying ophiolite could be similar or different and may provide
some constraint on the different models outlined in Fig. 1. Al-
though geochemical investigations of sole-ophiolite pairs may
prove useful, the age relationships discussed above can provide
further tests on the subduction initiation scenarios. Although all
four scenarios listed above predict similar ages for the initiation of
subduction (i.e., the metamorphic age of the sole) and the over-
lying ophiolite, each case predicts a different age relationship be-
tween the initially subducted crust (i.e., the metamorphic sole’s
protolith age) and subduction initiation. A determination of the
igneous crystallization age of the protolith of the high temperature
metamorphic sole has been the missing piece of information in all
previous studies of subophiolitic metamorphic soles and is key to
testing the various models of subduction initiation.
Cenozoic Subduction and Collision in Palawan
We applied the foregoing test to the metamorphic sole of an
ophiolite associated with a young, short-lived subduction zone in
Palawan, western Philippines (Fig. 2). The central Palawan
ophiolite was trapped in the forearc of the subduction zone that
generated the Cagayan arc (18). Subduction began at ∼34 Ma and
lasted for ∼20 My before it was terminated by microcontinent-arc
collision (18, 29, 33). The subduction zone’s relatively young age
older than both the
sole and the
protolith older than
the sole and slightly
older than the
protolith age similar to
from the sole and the
ophiolite ages Ages are compatible
with Model (D)
older than both the
sole and the
34.1 ± 0.1 Ma
Cooling of the
34.2 ± 0.6 Ma
of sole protoliths
35.25 ± 0.15 Ma
35.242 ± 0.062 Ma
35.862 ± 0.048 Ma
Fig. 1. Models of subduction initiation that explain similar ages between the formation of metamorphic soles and associated ophiolites (in cross-section and
map view). The high temperature metamorphic sole (shown as a thick, black line) is generated from the crust of the subducting plate during subduction
initiation. It may then be preserved at the base of the upper plate (future ophiolite, shown in cross-hatched pattern). Each model predicts a differentage
relation between the initially subducted crust, the overlying ophiolite, and the time of subduction initiation. Plate ages are schematically shown with darker
shades representing older lithosphere. White arrows on subducting plate indicate relative plate motion. (A) Sinking of the subducting plate along a transform
fault (TF) or fracture zone (FZ) drives extension in the upper plate, generating the future ophiolite (10). (B) Subduction initiation of distinctly older lithosphere
near an active spreading center (27, 28). (C) Subduction initiation along an oceanic detachment fault near a spreading center (12). (D) Subduction initiation
very close to or at a spreading center axis (13). Hacker et al. (25) proposed a variant of this in which subduction initiates across a transform or fracture zone
with underthrusting directed parallel to an active spreading center axis. Both options are shown in map view. (E) Schematic of the Palawan ophiolite, its
metamorphic sole, and the dated lensoid pods preserved within the sole. U-Pb zircon ages of the pods and ophiolite obtained from this study are displayed
along with the metamorphic cooling age of the sole (18).
www.pnas.org/cgi/doi/10.1073/pnas.1609999113 Keenan et al.
and short duration make it an ideal site to study subduction initi-
ation because it was terminated before a potentially complicated
history might have ensued.
The island of Palawan flanks the southern margin of the South
China Sea (Fig. 2). Its geology consists of ophiolite-related rocks
(31, 32) and continental crust [the North Palawan continental
terrane (NPCT)] that rifted from southeast China during the
opening of the South China Sea Basin (SCSB) (29, 34). The
Cagayan Ridge is located to the southeast of, and trends parallel
to, Palawan. It is composed of calc-alkaline volcanic rocks (33) and
is the volcanic arc associated with the subduction zone whose in-
ception is preserved on Palawan (18, 29, 33, 34). To the northwest
of central and southwest Palawan is the Palawan trough (Fig. 2), a
linear depression that, based on seismic reflection observations, is
interpreted as a downwarped segment of the southern edge of the
NPCT and contiguous proto-SCSB that is underthrust beneath the
Palawan ophiolite (35, 36). The proto-SCSB is Cretaceous oceanic
lithosphere that existed south of the NPCT–southeast China pas-
sive margin before opening of the SCSB (29). Allochthonous
remnants of this older Cretaceous ophiolite are found in tectonic
windows beneath the ∼34 My old (see below) Palawan ophiolite
(18, 31, 36). Gravity data are consistent with the younger Central
Palawan ophiolite being rooted in the south (37) and thrust
northward on to the continental crust of the NPCT (Fig. 2D)(32)
in a manner similar to Tethyan ophiolites (20).
In summary, all of the available evidence (onshore geologic and
offshore drill hole, seismic, and gravity data) from Palawan and
the surrounding areas is consistent with a south-southeast dipping
Basalt / Diabase
pillow basalts (L. Cretaceous
to E. Eocene)
Clastics and Carbonates
(Upper Eocene? - Miocene)
Alluvium/ River Deposits
(Pliocene - Recent)
oceanic crust of
NW limit of Palawan
NPCT Ulugan Bay
118 °0 0’ 11 8° 45 ’
with hornblendite lenses
Location of photos
in Figure 3
NPCT PAL AWA N
TROUGH PAL AWA N
Schematic NW-SE section across Ulugan Bay area (see panel “b”)
DERIVED FROM NPCT
Fig. 2. (A) Present tectonic setting of Palawan island, Philippines (18, 29, 30). Rectangle outlines the area shown in B.(B) General geology of central Palawan
showing locations of sample sites. The general structure consists of an ∼34-Ma ophiolite (the Central Palawan ophiolite) thrust over deformed Cretaceous-
Eocene turbiditic sedimentary rocks of the NPCT. Remnants of the older Early Cretaceous proto-SCSB ophiolite are found as occasional pillow lavas in tectonic
windows in the younger Palawan ophiolite (geology from our field observations and refs. 18 and 31). (C) Geologic map of the metamorphic sole at Dalrymple
Point. See Bfor location. Background image from Google Earth (Digital Globe, CNES/Astrium). Apparent metamorphic grade decreases away from the mantle
peridotite. (D) Schematic NW-SE cross section of Palawan in the Ulugan Bay area after ref. 32.
Keenan et al. PNAS Early Edition
AND PLANETARY SCIENCES
subduction zone that formed the Cagayan arc, which places the
Palawan ophiolite in the forearc. This subduction zone was ter-
minated when the southern edge of the rifted margin of China
(the NPCT) jammed the trench in the Middle Miocene (33),
causing obduction of the Palawan ophiolite. The question of how
subduction initiated in Palawan is unresolved. Mitchell et al. (32)
speculated that subduction may have initiated at a spreading
center, whereas the model of Encarnación et al. (18) shows initi-
ation at a transform fault or fracture zone. Both of these proposals
lack the necessary age data to properly test these models.
Age and Geochemistry of the Palawan Ophiolite and Its Sole
The central Palawan ophiolite is a relatively coherent section of
oceanic lithosphere consisting of pillowed and massive lavas, diabase
dikes, plagiogranite, gabbro, troctolite, and mantle harzburgite (31,
32) (Fig. 2). Samples of pillow lavas, dikes, and gabbroic intrusions,
as well as felsic intrusions (plagiogranite) (Fig. 3) collected from
the ophiolite, show slight light rare earth element (LREE) de-
pletions and relatively flat middle REE and heavy REE patterns,
about 10 times chondritic values, consistent with midocean ridge
(MORB)-like magma (Fig. 4A). A similar MORB to transitional
MORB-island arc basalt (IAB) or suprasubduction zone signature
is evident in other multielement plots. Tectonic discrimination
diagrams that use robust statistical tests (38) assign the data within
the MORB field (Fig. 4B) or to a transitional MORB-IAB field
(Fig. 4C), consistent with a backarc basin basalt (BABB)-type
geochemistry (39) and consistent with many other ophiolites (40)
formed at oceanic divergent-type boundaries.
Based on zircon U-Pb geochronology from a plagiogranite
sampled near Penacosa Point (Fig. 2), the best estimate for the
age of the ophiolite is 34.1 ±0.1 Ma (Fig. 1 and Table S1). The
regional distribution of rock types (Fig. 2) shows that the outcrops
of gabbroic and tonalitic intrusives in the Penacosa Point area are
located in the upper levels of the main gabbroic crustal section
transitional to the extrusive section of the ophiolite. The Penacosa
Point plagiogranite exhibits field relations that are consistent with
the felsic magma being comagmatic with the dominant mafic
magmas comprising the bulk of the oceanic crustal section here
(Fig. 3). In addition, the geochemistry of the plagiogranite is
consistent with simple fractional crystallization from the dominant
basaltic magma (Fig. 4A).
The metamorphic sole of the ophiolite is exposed in the Dal-
rymple Point area and is composed of ductiley strained garnet
amphibolites, hornblendites, amphibolites, epidote amphibolites,
quartzites, and kyanite schists with isoclinal folds and a distinct,
penetrative mineral lineation in most rocks (Fig. 2). These rocks
represent metamorphosed igneous, basaltic oceanic crust and as-
sociated sediments (cherts and mudstone). As in other meta-
morphic soles, the higher temperature garnet amphibolites and
hornblendites are located closer to the basal mantle harzburgites,
whereas lower temperature epidote amphibolites tend to be
structurally lower. At several locations, more highly strained sole
rocks enclose less deformed (or isotropic) and more competent,
irregularly shaped lensoid pods of epidote amphibolite (decimeter-
meter scale) and smaller lensoid pods of amphibolite (a few mil-
limeters to centimeters in thickness and several centimeters to
decimeters in diameter) (Fig. 3). Thermobarometric determina-
tions show that the garnet amphibolites of the sole reached peak
metamorphic temperatures of 700–760 °C and pressures exceeding
9kbar(∼27 km depth in the mantle) (18), conditions consistent
with those that the crust of a subducting plate would be subject to
during the earliest stages of subduction (23). Critical to this study,
previous work (18) has constrained the minimum age for sub-
duction initiation by determining
Ar cooling ages on two
hornblende samples and one white mica sample (from garnet
amphibolite, amphibolite, and kyanite schist, respectively) in the
sole. These ages are indistinguishable at 34.2 ±0.5, 34.2 ±0.6, and
34.25 ±0.3 Ma, respectively (Fig. 1) (corrected for new revised
ages of the neutron flux monitors) (41), and indicate rapid cooling
of the sole to 550–400 °C after reaching peak metamorphic tem-
peratures (18). Lower grade, altered greenstones, and pillow lavas
structurally beneath the metamorphic sole are exposed further
south in the Sagasa Point area (Figs. 2 and 3). These rocks are less
metamorphosed oceanic components underthrust beneath the
Palawan ophiolite and its sole sometime after the initial un-
derthrusting associated with subduction initiation.
Overall, metabasite samples from the metamorphic sole are
geochemically similar to the ophiolite in that they plot in the
MORB-like to transitional MORB-IAB fields (Fig. 4). Several
samples of the epidote amphibolite pods are depleted (∼2–3 times
chondritic values) relative to the MORB-like samples and have
positive Eu anomalies, indicating they were probably cumulate
50 um 100 um 100 um 100 um 100 um
GH I J K L
Fig. 3. Photographs of outcrops from the central Palawan ophiolite (A) and its metamorphic sole (B–F) and cathodoluminescence images of extracted zircons
from selected samples (G–L). (A) Magma-mingling structures exhibited by light-colored tonalite (plagiogranite) and diorite-gabbro (darker) at Penacosa Point.
The tonalite yielded zircons with a crystallization age of 34 Ma. Pencil for scale. (B) Layered chert/quartzite and amphibolite showing sheath folds.
(C) Amphibolite gneiss with hornblendite domains exhibiting isoclinal folding. (D) Foliated and lineated epidote amphibolite, looking ∼west; mountains
across Ulugan Bay are mantle harzburgite of the Palawan ophiolite structurally overlying metamorphic sole rocks; strike and dip symbol indicates foliation.
(E) Smaller, foliation-parallel, light-colored lensoid pods of amphibolite (sample PL-14-07). (F) Competent, light-colored pod of epidote amphibolite (with
cumulate gabbro-like REE signatures; sample PL-14-05) enclosed within the strongly foliated amphibolite. These pods yielded zircons with crystallization ages
of 35.242 Ma. Hammer for scale. (Gand H,Iand J, and Kand L) Cathodoluminescence images of zircons, showing magmatic oscillatory zoning (samples PL-14-
05, PL-14-06, and PL-14-07, respectively).
www.pnas.org/cgi/doi/10.1073/pnas.1609999113 Keenan et al.
rocks (gabbros, based on major element chemistry) of the mid-
lower crust. One of the smaller pods of amphibolite displays a
negative Eu anomaly and is slightly more enriched than the larger
epidote amphibolite pods. This sample may have crystallized from
a magma following the extraction of plagioclase. The rocks of the
sole are therefore not geochemically unlike the overlying ophiolite
and probably formed in a similar petro-tectonic setting.
Based on their geochemistry and geologic context within the
high temperature sole, we interpret the more competent pods to
be middle- to lower-level crustal rocks of the leading edge of the
subducted plate. During the period of underthrusting to sub-
duction, the underthrust crust was sheared, thinned, and ductiley
folded, resulting in the transposition of middle to lower crustal
gabbroic rocks with upper crustal basaltic rocks, both of which
were subject to the high temperature and high pressure meta-
morphism that formed the metamorphic sole (Fig. 1).
Zircons from two of the larger pods (PL-14-05 and PL-14-06)
and one smaller pod (PL-14-07) from the high temperature
chronology. The four analyzed zircons from sample PL-14-05
yielded internally and externally concordant ages with an error-
weighted mean age of 35.242 ±0.062. Three of five analyzed
zircons from sample PL-14-06 yielded internally and externally
concordant ages with a mean of 35.862 ±0.048 Ma (Fig. 1, Table
S2,andFig. S1). Three externally discordant but internally con-
cordantagesof37.00±0.16, 35.97 ±0.11, and 35.25 ±0.15 Ma
(Fig. 1, Table S2,andFig. S1) were obtained from the smaller pod.
The age of the youngest zircon, 35.25 ±0.15 Ma, is taken as the
best estimate of the final crystallization age of the protolith of this
sample. The small spread in ages seen in this sample is similar to
those revealed by high precision U-Pb geochronology in the
Samail ophiolite (44) and may be due to prolonged zircon crys-
tallization in a replenished magma chamber or assimilation of
slightly older wall rock.
Establishing that these zircons are igneous, and not meta-
morphic, is critical because the age of metamorphic zircons would
merely represent the age of metamorphic sole formation (sub-
duction initiation) instead of the crystallization age of the meta-
morphic sole protoliths. Although metamorphic zircon growth has
been shown to occur under amphibolite facies conditions (45),
cathodoluminescence imaging shows no evidence of metamorphic
overgrowths in these zircons. Instead, they are euhedral, prismatic,
and have distinct, fine, oscillatory zoning, a feature that is char-
acteristic and unique to magmatic zircons (46) (Fig. 3). Even
though Th/U ratios are not completely reliable indicators of
magmatic vs. metamorphic zircon, we note that the Th/U ratios in
these zircons (>0.1) are consistent with many magmatic zircons
(47). We are therefore confident that these zircons are igneous and
that their ages reflect the original igneous crystallization age of the
oceanic crust that was underthrust, and then metamorphosed, at
the onset of subduction.
Forced Subduction of Young, Buoyant Lithosphere
As discussed earlier, a critical test to constrain the tectonic setting
of subduction initiation is a comparison of the igneous ages of the
underthrust and overriding lithosphere in relation to the time of
subduction initiation. A positive test for subduction initiation at an
active spreading center is to find all three events very close in age.
We find that the age differences between the upper plate (the
Palawan ophiolite), the subducting plate (protoliths of the sole),
and metamorphism of the sole are less than ∼1My(Fig.1).The
very small age difference between formation of the sole protolith
and its metamorphism during subduction initiation leads us to re-
ject outright the models shown in Fig. 1 Aand B.Furthermore,the
model shown in Fig. 1Cis rejected for subduction initiation at
detachment faults that are far from the spreading center. We
therefore conclude that subduction must have initiated in very
close proximity to, or at, a spreading center (Fig. 1D). Our age data
do not differentiate between the ridge parallel and ridge normal
subduction initiation scenarios in Fig. 1Din cases where the ridge-
transform fault segments in the right-hand scenario are very short.
The similarity in the geochemistry of the sole and overlying
ophiolite also supports subduction initiation close to a spreading
center, where the eventual upper plate and lower plate (meta-
morphic sole) are not expected to be geochemically different.
Our data show that oceanic crust was formed at 35.24 Ma and
was then underthrust/subducted almost immediately, reaching
∼27 km depth, metamorphosed to amphibolite, and subsequently
cooled to ∼400 °C by 34.25 Ma, a remarkably short interval be-
tween crust formation and subduction. Assuming a slab dip con-
trolled by an ∼30° dipping lithosphere–asthenosphere boundary,
this underthrusting would require a convergence rate on the order
of 5 cm/y.
In cases where the detachment fault (Fig. 1C)islocatedvery
close to, or at the spreading center, the models shown in Fig. 1 C
and Dbecome indistinguishable using age data alone. The weak-
ness at which subduction initiated could have been a detachment
fault very near the spreading center axis or at the spreading center
axis itself. Although the exact nature of the weak zone could be
debated, the high-precision age data from our sample site tightly
- Metamorphic Sole
- Palawan ophiolite
Ti / 50
Ti / 50
V50 * Sm 5 * Sc
Sample / Chondrite
- Metamorphic Sole (dated samples)
- Metamorphic Sole
- Palawan ophiolite (dated sample)
- Palawan ophiolite PW-00-18
Fig. 4. Geochemical data on samples from the central Palawan ophiolite
and its metamorphic sole. Palawan ophiolite samples plotted in A–Cinclude
pillow lavas, mafic dikes, gabbroic intrusions, and (A) felsic intrusions (pla-
giogranite). Metamorphic sole samples plotted in A–Cinclude amphibolites,
epidote amphibolites, and garnet amphibolites. (A) Chondrite normalized
REE concentrations in the samples. Samples that were selected for U-Pb
zircon geochronology are symbolized by diamonds and their ages are in-
dicated next to the data. The majority of samples have REE patterns re-
sembling MORB and possible differentiates of MORB. Two samples (PL-14-06
and PL-14-05) display positive Eu anomalies indicating cumulate plagioclase
in the samples. The geochemistry of the plagiogranite (PW-00-18) is consis-
tent with simple fractional crystallization from the MORB-like basaltic
magmas. (B) Ti-V-Sm and (C) Ti-V-Sc tectonic discrimination diagrams (38).
Basaltic samples from the ophiolite and sole are similar and plot as MORB or
Keenan et al. PNAS Early Edition
AND PLANETARY SCIENCES
constrains subduction initiation at, or very close to the spreading
axis. If a detachment fault was the weakness, it would have to have
been located very close to the ridge where nascent oceanic crust
was generated and then underthrust and metamorphosed at a
depth of ∼27 km less than 1 My later.
The much younger age of the Palawan ophiolite compared with
the Cretaceous proto-SCSB suggests that subduction was initiated
within a young marginal oceanic basin hosted in older Cretaceous
proto-SCSB (Fig. 5). Because of the positive buoyancy of young
oceanic lithosphere, subduction initiation must have been forced
in this case. Following forced underthrusting of the young, buoy-
ant lithosphere, the old, cold, and dense Cretaceous lithosphere of
the proto-SCSB would have eventually entered the trench and
enabled the transition to self-sustained subduction until the NPCT
collided with the trench, caused obduction of the Palawan
ophiolite and its sole, and terminated subduction.
Regionally, the origin of the force that converted a divergent
boundary to a convergent boundary was probably the collision of
India with Asia. Tapponnier et al. (49) proposed that this collision
resulted in the extrusion of fairly rigid continental lithospheric
blocks from the southeast margin of Asia along large strike-slip
faults, such as the Red River shear zone (Fig. 2). The timing of the
onset of strike-slip movement along the Red River shear zone
has been estimated by U-Pb ages on monazite included in shear-
related rotated garnets at ∼34 Ma (48). This age coincides well
with the timing of the initiation of subduction in Palawan. Seafloor
spreading in the South China Sea, which accompanied the south-
ward motion of the NPCT, began around 32–30 Ma (30), a time
also compatible with initiation of subduction in Palawan south of
the NPCT (18) and southward convergence of the NPCT with the
Palawan subduction zone–Cagayan arc.
If the reconstruction of Fig. 5 is correct, the initiation of sub-
duction at the Palawan spreading center, forced by the extrusion of
Indochina, supports the results of numerical modeling (1), which
predict that the thin lithosphere at oceanic spreading centers re-
quires less force to converge comparedwithareasof thicker lith-
osphere. Several potential weak zones existed in the area: (i)the
southeast China passive margin, (ii) the contact between the older
Cretaceous proto-SCSB lithosphere and the much younger Pala-
wan marginal basin, and (iii) the spreading center of the Palawan
ophiolite. Despite the presence of older and denser oceanic litho-
sphere, our results indicate that subduction nucleated within the
young, buoyant, but weaker Palawan marginal basin. Presumably,
after the eventual underthrusting of the older Cretaceous litho-
sphere and an additional critical convergence of 100–130 km,
subduction became self-sustaining (1).
There is currently no known modern analog for subduction
initiation exploiting an active oceanic spreading center. Intra-
oceanic subduction appears to be initiating along several sections
of the Australian–Pacific plate boundary south of New Zealand,
exploiting weaknesses associated with extinct spreading centers
and/or fracture zones undergoing transpressional deformation (50–
52). However, unlike the case in Palawan where deep un-
derthrusting occurred within 1 My of oceanic crust formation,
spreading at the boundary south of New Zealand had ceased and
was followed by strike-slip deformation several million years before
Cretaceous oceanic lithosphere
Detachment of Palawan ophiolite at
ridge and initiation of subduction
Onset of Red River
shear zone (~34 Ma)
Seafloor spreading starts in
South China Sea at 32-30 Ma
Detachment of Palawan ophiolite
at the spreading center and forced
initiation of subduction (~35-34 Ma)
due to extrusion of Indochina
Passive margin Extinct fault or
Fig. 5. Subduction initiation (∼34–35 Ma) at the spreading center that generated the Palawan ophiolite. (A) Schematic map view of the area showing the
timing of initial strike-slip movement on the Red River shear zone (48), seafloor spreading in the South China Sea (30), and initiation of subduction atthe
Palawan ophiolite spreading center. Palawan ophiolite (yet to be obducted onto the rifted continental crust) shown in cross-hatched pattern. (B) Cross-
sectional schematic view of the area along A-A′during the onset of subduction. Weak zones where subduction had the potential to initiate, but did not, are
www.pnas.org/cgi/doi/10.1073/pnas.1609999113 Keenan et al.
incipient subduction. It should be noted that there are only a few
locations where subduction initiation might be happening (ref. 1
and references therein), and therefore it is unlikely that these
represent the full range of potential mechanisms for intraoceanic
subduction initiation. The process of subduction initiation may last
on the order of only 2–3 My (assuming a convergence rate of
∼5 cm/y) (1), a small fraction of the lifetime of a subduction zone,
and is therefore unlikely to be captured by present day observa-
tions. Ophiolites that have similar crystallization ages to the
metamorphic age of their sole, like the Palawan ophiolite, are not
uncommon (19, 25, 26). These other examples, however, lack dates
for the protolith of the sole. Dating these protoliths may show that
subduction initiation at spreading centers has been more pervasive
than currently recognized.
The subduction zone that existed in Palawan and the nascent
Puysegur trench (50) are possibly the only well-constrained geo-
logic examples that address the dynamics of subduction initiation.
In both of these examples, a forced initiation has been inferred.
Although spontaneous subduction has been simulated by numer-
ical models (9), corresponding geologic tests have been questioned
(53). Our study suggests that subduction initiation in the modern
day plate tectonic regime requires extant subduction, the main
driver of plate motion, to force new subduction. Furthermore, it
appears that forces generated in the interior of major continental
collisions zones can be transmitted out to the oceanic realm,
rapidly causing diverging plates at a spreading center to converge
and lead to incipient subduction in less than ∼1My.
Geochemical Analysis. Whole rock powders were analyzed by a combination of
a lithium metaborate/tetraborate fusion followed by inductively coupled
plasma (ICP) methods for major elements abundances and inductively coupled
plasma MS(ICP-MS) methods fortrace elements abundances.The analyses were
done at Activation Laboratories (Ontario, Canada) (ACTLAB). A detailed de-
scription of sample preparation methods is given on the ACTLAB Website
U-Pb Zircon Geochronology. All reported ages are
mean ages, because the uncertainty of
Pb ages are pronouncedly large
due to the very young and low
U content of these zircons. They are cor-
rected for initial Th/U disequilibrium, and errors are reported as 2σ.
Twelve zircon grains from the epidote amphibolite pods (PL-14-05, PL-14-06,
and PL-14-07) were analyzed by chemical abrasion thermal ionization MS
(CA-TIMS) at the radiogenic isotope laboratory at Massachusetts Institute of
Technology. Samples PL-14-05 and PL-14-06 yielded weighted mean ages of
35.242 ±0.062 and 35.862 ±0.048 Ma, respectively, whereas sample PL-14-07
displays three distinct ages of 37.00 ±0.16, 35.25 ±0.15, and 35.97 ±0.11 Ma.
For sample PL-14-06, the youngest cluster of zircon ages was used for the
weighted mean age because this best approximates the timing of magma
Three zircon fractions from the plagiogranite (PW-00-18) were analyzed by
conventional TIMS at the University of California, Santa Barbara. One fraction
was analyzed as is, whereas the other two were air abraded to remove any
exterior zones with possible Pb-loss. The unabraded fraction is only slightly
younger thanthe two abraded fractions that give a mean age of 34.1 ±0.1 Ma.
Additional zircon fractions were analyzed using the stepwise chemical abra-
sion technique (55), and all analyzed fractions after the first few steps yielded
identical 34.1 ±0.1 Ma ages.
ACKNOWLEDGMENTS. We thank John Spray and two anonymous reviewers
for helpful and constructive comments. Freddie Dela Cruz provided excellent
boatmanship in Ulugan Bay. This work was supported, in part, by Saint Louis
University and the Geological Society of America. The geochronology
laboratories at University of California Santa Barbara and Massachusetts
Institute of Technology are supported by the National Science Foundation.
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Keenan et al. 10.1073/pnas.1609999113
0.00554 0.00556 0.00560 0.00562
0.0356 0.0358 0.0360
weighted mean 206Pb/238U date
35.242 ± 0.046/0.050/0.062 Ma
MSWD = 0.72, n=4
weighted mean 206Pb/238U date
35.862 ± 0.024/0.029/0.048 Ma
MSWD = 0.86, n=3
0.034 0.035 0.036 0.037 0.038 0.039 0.0400.033
0.0054 0.0055 0.0057 0.00580.0056
Fig. S1. U-Pb Concordia diagram for samples from the metamorphic sole. (A) Data from PL-14-05. Weighted mean age is reported as 35.242 ±0.062 Ma.
(B) Data from PL-14-06. Weighted mean age is reported as 35.862 ±0.048 Ma. (C) Data from PL-14-07. Three externally discordant, but internally concordant,
ages are 37.00 ±0.16, 35.97 ±0.11, and 35.25 ±0.15 Ma.
Keenan et al. www.pnas.org/cgi/content/short/1609999113 1of3
Table S1. Zircon U-Pb analytical results from plagiogranite (PW-00-18) of Central Palawan Ophiolite
Isotopic ratios Ages
100–140 abr 1 0.2796 13.54 60.19 0.20051 0.05624 1362.4 0.005302 34.09
70–100 abr 0.5 0.2575 6.66 55.41 0.20742 0.0575 1137.7 0.005301 34.08
70–100 uabr 4 0.5382 65.00 117.4 0.19526 0.0528 2409.6 0.005255 33.79
Zircon sizes are in micrometers. Uncertainties on the
U ages are ±0.2% (2σ). Analyses were done by thermal ionization MS at the University of
California, Santa Barbara. abr, abraded; uabr, unabraded.
Keenan et al. www.pnas.org/cgi/content/short/1609999113 2of3
Table S2. Zircon U-Pb analytical results from epidote-amphibolite pods (PL-14-05 and PL-14-06) and amphibolite pod (PL-14-07) of the metamorphic sole
Composition Isotopic ratios Dates (Ma)
Z2 1.16 0.3 15.81 817.4 0.373 0.005474 0.79 0.03570 2.05 0.047316 1.817 35.19 0.28 35.61 0.72 64 43 0.47
Z4 1.16 0.2 13.86 719.8 0.371 0.005485 0.21 0.03561 2.07 0.047109 1.986 35.262 0.074 35.53 0.72 54 47 0.44
Z5 0.65 0.3 17.19 1001.2 0.209 0.005476 0.21 0.03532 1.76 0.046801 1.629 35.205 0.072 35.25 0.61 38 39 0.69
Z6 0.90 0.6 6.95 391.4 0.290 0.005488 0.29 0.03573 3.75 0.047234 3.648 35.28 0.10 35.6 1.3 60 87 0.39
Z1 0.23 0.4 35.9 2321.8 0.074 0.005577 0.12 0.03591 0.65 0.046712 0.618 35.855 0.042 35.82 0.23 33 15 0.36
Z2 0.25 0.3 25.52 1645.5 0.081 0.005582 0.11 0.03605 0.89 0.046867 0.858 35.881 0.038 35.96 0.32 41 21 0.38
Z3 0.20 0.3 35.5 2317.6 0.064 0.005602 0.15 0.03613 0.72 0.046800 0.666 36.014 0.055 36.04 0.25 38 16 0.42
Z4 0.22 0.3 22.84 1490.0 0.069 0.005576 0.13 0.03605 1.11 0.046918 1.064 35.844 0.047 35.96 0.39 44 25 0.38
Z5 0.25 0.7 25.86 1667.5 0.080 0.005592 0.14 0.03633 0.96 0.047134 0.894 35.950 0.050 36.23 0.34 55 21 0.51
Z2 0.53 0.4 4.80 301.6 0.169 0.005756 0.44 0.03726 5.44 0.046969 5.229 37.00 0.16 37.1 2.0 47 125 0.50
Z3 0.44 0.5 3.90 254.0 0.142 0.005483 0.43 0.03543 5.63 0.046884 5.487 35.25 0.15 35.4 2.0 42 131 0.37
Z4 0.41 0.4 6.28 400.7 0.132 0.005596 0.30 0.03650 3.59 0.047328 3.466 35.97 0.11 36.4 1.3 65 83 0.43
Pb =18.15 ±0.48;
Pb =15.30 ±0.29;
Pb =37.11 ±0.88; Mass fractionation correction of 0.25%/amu ±0.02%/amu (atomic mass unit) was applied to all single-
collector Daly analyses.
Th contents calculated from radiogenic
Pb and the
Pb date of the sample, assuming concordance between U-Th and Pb systems.
Total mass of common Pb.
Ratio of radiogenic Pb (including
Pb) to common Pb.
Measured ratio corrected for fractionation and spike contribution only.
Measured ratios corrected for fractionation, tracer, and blank. All common Pb was assumed to be procedural blank. Total procedural blank for U was less than 0.1 pg.
Corrected for initial Th/U disequilibrium using radiogenic
Pb and Th/U [magma] =2.8.
**Isotopic dates calculated using the decay constants of λ
(42), and for the
U=137.818 ±0.045 (43).
Keenan et al. www.pnas.org/cgi/content/short/1609999113 3of3