# A coral microatoll record of sea-level rise, interseismic deformation, and El Niño in La Union, Philippines since 1906 CE

Gina M Sarkawi, Aron J Meltzner, Dongju Peng, Joanne T Y Lim, Xinnan Li, Rohan Gautam, Anandh Gopal, Andrew Mitchell, Loraine Faye Sarmiento, Jennifer Weil-Accardo, et al.

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## To cite this version:

Gina M Sarkawi, Aron J Meltzner, Dongju Peng, Joanne T Y Lim, Xinnan Li, et al.. A coral microatoll record of sea-level rise, interseismic deformation, and El Niño in La Union, Philippines since 1906 CE. Marine Geology, 2025, 486, pp.107565. 10.1016/j.margeo.2025.107565. hal-05450990

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Research Article

# A coral microatoll record of sea-level rise, interseismic deformation, and El Niño in La Union, Philippines since 1906 CE

Gina M. Sarkawi<sup>a,b,*</sup>, Aron J. Meltzner<sup>a,b</sup>, Dongju Peng<sup>a,c</sup>, Joanne T.Y. Lim<sup>a,b</sup>, Xinnan Li<sup>a,d</sup>, Rohan Gautam<sup>a,b</sup>, Anandh Gopal<sup>a,b</sup>, Andrew Mitchell<sup>a,b,e</sup>, Loraine Faye Sarmiento<sup>f,g</sup>, Jennifer Weil-Accardo<sup>a,h</sup>, Kathrine V. Maxwell<sup>a,i</sup>, Junki Komori<sup>a,b</sup>, Noelynna T. Ramos<sup>f</sup>

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ARTICLE INFO

Editor: Shu Gao

Keywords:

Coral microatolls

El Niño–Southern Oscillation

Relative sea level

Philippines

ABSTRACT

Relative sea level (RSL) along western Luzon is influenced by climate change, tectonic deformation, and the El Niño–Southern Oscillation. Long, high-resolution records of past changes in sea level and land level are crucial to distinguish these contributions for forecasting future RSL or identifying tectonic hazards. Coral microatolls are invaluable proxies for reconstructing RSL, but their interpretation relies on understanding how short- and long-term changes are recorded in the coral skeleton. We analyzed growth bands from 1906 to 2018 CE of a *Porites* microatoll from Balaoan, La Union, Philippines, using X-rays of vertical cross sections. The microatoll's surface morphology, annual banding, and growth history reveal RSL rise during its lifetime punctuated by 17 diedowns. Coral diedowns generally coincide with low water near the site, often resulting from El Niño events. Excluding eroded sections, 10 of the 11 strongest El Niños since 1918 produced a diedown. Similarly, excluding a diedown caused by the coral tilting, 10 of the 12 preserved diedowns exceeding 3 mm in amplitude occurred during strong El Niños. Between 1948 and 2018 CE, RSL rose at 1.3 ± 1.8 mm/yr (2σ). Comparing sea-level rates from satellite altimetry and RSL from the coral, we estimate vertical land motion at −0.1 ± 12.1 mm/yr between 1995 and 2018 CE. The indicative meaning at Balaoan, defined as the relationship between local water levels and the highest level of growth of a coral microatoll before its most recent diedown, is estimated at 31.5 ± 10.0 cm (2σ) above lowest astronomical tide.

# 1. Introduction

Sea-level rise is an ongoing concern for coastal communities, but quantifying current rates of sea-level change and forecasting future sea level are hindered by the fact that sea level itself is complicated in space and time. Globally averaged rates of sea-level rise, 1.1–2.0 mm/yr since the 20th century (Dangendorf et al., 2017; Frederikse et al., 2020; Hay et al., 2015; Palmer et al., 2021), do not capture the regional variability of relative sea-level (RSL) change. RSL is a function of changes in both sea-surface height (SSH) and land level (e.g., Stammer et al., 2013).

Ocean–atmosphere interactions such as the El Niño–Southern Oscilla-
tion (e.g., Cazenave et al., 2012) or the Pacific Decadal Oscillation (Chen
et al., 2020) affect SSH on interannual to decadal scales. Land-height
changes in the form of tectonic deformation (e.g., Montillet et al.,
2018), glacial isostatic adjustment (GIA; e.g. Mitrovica and Milne,
2002), or localized subsidence due to compaction or groundwater
withdrawal (e.g., Shirzaei et al., 2021) are also recorded in the RSL
signal. As a result of this complexity and the sparsity of long-term tide-
gauge records in tropical regions (Church and White, 2011), RSL
changes are particularly poorly quantified in lower latitudes.

<footnote><sup>a</sup> Earth Observatory of Singapore, Nanyang Technological University, Singapore</footnote>

<footnote><sup>b</sup> Asian School of the Environment, Nanyang Technological University, Singapore</footnote>

<footnote><sup>c</sup> Department of Land Surveying and Geo-Informatics, The Hong Kong Polytechnic University, Hong Kong</footnote>

<footnote><sup>d</sup> State Key Laboratory of Earthquake Dynamics, Institute of Geology, China Earthquake Administration, Beijing, China</footnote>

<footnote><sup>e</sup> Lancaster Environment Centre, Lancaster University, Lancaster, United Kingdom</footnote>

<footnote><sup>f</sup> National Institute of Geological Sciences, College of Science, University of the Philippines, Diliman, Quezon City, Philippines</footnote>

<footnote><sup>g</sup> Global Geologica Pty Ltd., Philippines</footnote>

<footnote><sup>h</sup> Aix Marseille University, CNRS, IRD, INRAE, College de France, CEREGE, Aix-en-Provence, France</footnote>

<footnote><sup>i</sup> Senckenberg am Meer Wilhelmshaven, Marine Research Department, Wilhelmshaven, Germany</footnote>

<footnote>* Corresponding author at: School of Earth and Environment, Universtiy of Canterbury, Christchurch, New Zealand.
E-mail address: gina.sarkawi@canterbury.ac.nz (G.M. Sarkawi).</footnote>
Coral microatolls are a viable proxy in the tropics to address short-
comings in our understanding of sea level, allowing for RSL re-
constructions in regions that otherwise lack information and by
extending sea-level records back into the pre-instrumental era (Meltzner
et al., 2010; Taylor et al., 1987; Weil-Accardo et al., 2016). Microatolls
form as a colony of coral polyps that grow in shallow intertidal waters,
and their vertical growth is limited by low water level (Meltzner and
Woodroffe, 2015; Scoffin and Stoddart, 1978). Changes in the water
depth, due to changes in either SSH or land level, influence the growth
pattern of the microatoll. The surface morphology of a coral microatoll is
characterized by concentric rings that form as a result of a temporary
low-water event, as subaerial exposure causes the uppermost surface of
the coral to die; such an event is termed a diedown, and its amplitude
records the magnitude of the RSL change (Meltzner and Woodroffe,
2015). These corals also form growth bands annually that record the
precise timing (to within a fraction of a year) of each diedown event
(Meltzner and Woodroffe, 2015). In a region of rising RSL, the surface
morphology of a coral microatoll would have outer rings at higher ele-
vations than its inner rings (Hopley, 1982). Where RSL is falling, there is
a net decrease in elevation from the center of the coral microatoll to its
outer rings (Hopley, 1982). Rates of RSL change can be determined
based on time series reconstructed from elevations and ages of different
parts of the coral microatoll.

Coral microatoll studies have proven useful in reconstructing RSL in
a variety of tectonic and climatic settings, including active subduction
zones (e.g., Sumatra; Meltzner et al., 2010), stable mid-oceanic islands
(e.g., the Cocos Islands; Smithers and Woodroffe, 2001), and areas
influenced by the El Niño–Southern Oscillation (e.g., Christmas Island;
Woodroffe and Gagan, 2000). These applications demonstrate the
versatility of microatolls as archives for reconstructing long-term and
abrupt changes in RSL, particularly in data-sparse regions. As such, this
approach is well suited for other low-latitude coastlines that lack long
instrumental records but experience significant tectonic or oceano-
graphic variability.

However, while a general correlation has been observed between the
timing of coral diedowns and periods of low water, the specific causes
and conditions behind these events are not well understood. This lack of
clarity limits the precision with which we can use coral microatolls to
track secular change in RSL. To overcome this problem, some studies
have attempted to compare independent sea-level time series to the
timings of coral microatoll diedowns (Mann et al., 2016; Meltzner et al.,
2017; Meltzner et al., 2010; Philibosian et al., 2022; Weil-Accardo et al.,
2016; Woodroffe and McLean, 1990; Zachariasen et al., 2000). The only
prior study that directly compares diedowns to a long-term hourly tide-
gauge record is that by Majewski et al. (2022). Their study compared
corals at an island in Indonesia with tide-gauge records in Singapore, but
their comparison was limited by the >100 km distance between the
corals and the tide gauges.

Here we present a RSL reconstruction from a coral microatoll in
western Luzon, Philippines, and compare it to an hourly intermittent
tide-gauge record from 1988 to 2021 located ~22 km away, offering the
most proximal coral microatoll slab to hourly tide-gauge comparison yet
published. Due to temporal gaps in the tide-gauge record, we reconstruct
hourly water levels by combining a tide model constructed from the
tide-gauge record with satellite altimeter-derived sea-level anomalies to
reconstruct water levels at our site. Reconstructed water levels provide a
continuous tide record with which to compare coral microatolls. To
reinforce the comparison between coral diedowns and low-water events,
we also compare the timing of coral diedowns with a standardized index
of the El Niño–Southern Oscillation (Webb and Magi, 2022), as El Niño
events tend to result in lower sea levels along the west coast of Luzon (e.
g., Peng et al., 2013). In addition to providing new insight into the
relationship between coral microatoll growth, low-water events, and
longer-term trends in RSL, our study offers insight into the drivers of RSL
change in western Luzon. We attempt to discretize the contributions of
land-level and sea-level change by quantifying vertical land motion

(VLM) using paired RSL and satellite-altimetry sea-surface data.

## 2. Study site and local drivers of RSL

A population of living coral microatolls was found in the intertidal
zone of Barangay Almeida of Balaoan municipality, La Union province,
hereafter referred to as Almeida (Fig. 1). The nearest tide-gauge station
is located ~22 km south-southwest in San Fernando, La Union (Fig. 1A).
Diurnal tides dominate this region, and the tidal range was quantified as
0.65 m between mean low water and mean high water based on the San
Fernando tide gauge (Supporting Table S1). Both living and fossil
microatolls were found in a partially open lagoon at the Almeida site
(Fig. 1B), behind an outer reef rampart. From aerial imagery (Support-
ing Fig. S1), the lagoon is not fully enclosed; an opening ensures con-
nectivity to the ocean currents. A small stream flows into the lagoon,
depositing mostly sand in a small delta; longshore drift transports some
of the sand southward along a sandy beach, and much of this sand fills
the deeper parts of the lagoon at the southern end of the site.

SSH at the site can be influenced on short timescales, over the course
of several weeks to months, during two known atmospheric phenomena.
Variability in surface winds pushes water to or from a coastline resulting
in changes in SSH. The El Niño–Southern Oscillation (e.g., Peng et al.,
2013) and the East Asian monsoon (e.g., Wang et al., 2004) are two
atmospheric phenomena in the region that produce surface wind vari-
ability on an inter- or intra-annual period. In the central to western
Pacific, east of the Philippines, SSH is temporarily lowered during ca-
nonical El Niño events. Still, models suggest this signal is weaker (Chang
et al., 2013; Landerer et al., 2008), more complex (Pei et al., 2015), or
both in the South China Sea.

Tectonic deformation may also influence RSL on short or long
timescales through land-level changes. The Manila subduction zone lies
~132 km west of our site. At the Manila trench, the Sunda plate con-
verges with the Philippine mobile belt at a rate of 83–94 mm/yr (Hsu
et al., 2016; Kreemer et al., 2003; Yu et al., 1999). The strain from
oblique convergence of the Sunda plate to the Philippine mobile belt is
partially partitioned between subduction along the Manila trench and
sinistral strike-slip along the Philippine Fault system onshore (Aurelio,
2000; Barrier et al., 1991). The largest recorded event along the Manila
subduction zone in the 20th century occurred in 1934. It was a M~s~ 7.6
located approximately 127 km offshore Vigan, Ilocos Sur, although its
exact location is not well constrained (Gutenberg and Richter, 1954;
Hamburger et al., 1983). The event produced a tsunami with reports of a
wave in San Esteban and surges in Vigan (Berninghausen, 1969). RSL
reconstructions along the west coast of Luzon (e.g., Maeda et al., 2004;
Maxwell et al., 2018; Ramos et al., 2017; Ramos and Tsutsumi, 2010)
have primarily focused on determining long-term tectonic deformation
in the region, but it is unclear how tectonics may influence local RSL at
different stages of the seismic cycle. Early Holocene records of RSL are
present from Balaoan based on cores from the uplifted coral reef margin
of T2 (Fig. 1B; Gong et al., 2013; Siringan et al., 2016). These records
indicate high levels of RSL that are consistent with paleosea-level re-
cords from further north in Luzon, at Currimao (Shen et al., 2010).

GIA, which describes the redistribution of water over gla-
cial-interglacial cycles and the viscoelastic response of the solid Earth to
the redistribution of that load, influences both SSH and land height, but
these are gradual processes that operate over long timescales. In low
latitudes, GIA is dominated by equatorial ocean siphoning, the migra-
tion of water in the far-field region towards the collapsed forebulge of
the previously glaciated region, and by continental levering, which
produces a spatially complex pattern along coastlines in which the
increased ocean load following deglaciation locally depresses the ocean
floor while uplifting areas farther inland (Khan et al., 2015; Mitrovica
and Milne, 2002).

Coastal subsidence can also occur, most significantly in river deltas,
due to natural compaction of the delta sediments and due to anthro-
pogenic groundwater extraction. Accelerated rates of present-day RSL
<figure><img src="image_5.png" /><figcaption>**Fig. 1.** (A) Regional map of western Luzon. Our study site is marked by the yellow square. The tide gauge (cyan) is located in San Fernando. Faults (black) on the main map are from Philippine Institute of Volcanology and Seismology (2023) while faults (red) in the inset map are based on Hinschberger et al., (2005), Meltzner et al. (2012), Mendoza et al. (2022), and Philippine Institute of Volcanology and Seismology (2023) (B) Site map showing living coral microatolls surveyed and slabbed at Balaoan, La Union, Philippines. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>

rise identified across populated coastal cities, including Manila, Jakarta, Ho Chi Minh City, and Chittagong, have been attributed to increased groundwater extraction and sediment compaction due to dense building load (e.g. Shirzaei et al., 2021; Tay et al., 2022).

Last but not least, eustatic sea-level rise due to climate change increasingly influences RSL at interannual to decadal timescales. Ongoing global temperature rise causes ocean mass increase due to melting of terrestrial glaciers (e.g., DeConto and Pollard, 2016), and ocean volume expansion due to the rise in ocean temperatures (e.g., Frederikse et al., 2020). This is an important problem that has received substantial global attention, but due to all the factors described above, it is impossible to resolve rates or amounts of eustatic sea-level change from studies at one site in isolation.

### 3. Coral microatolls as a proxy for relative sea level

A coral microatoll forms as a colony of coral polyps whose upward growth is controlled by extreme low-water (ELW) levels (Meltzner and Woodroffe, 2015). Microatolls tend to form on a shallow substrate and produce annual density bands which grow radially upwards and outwards (Scoffin and Stoddart, 1978). As the coral grows higher, or if the land rises or sea level falls, the coral is increasingly prone to periods of prolonged subaerial exposure at extreme low tide, which can kill the uppermost portion of the coral colony (Meltzner and Woodroffe, 2015). These diedowns are generally uniform in elevation around the perimeter of a microatoll, and portions of the microatoll that remain sufficiently submerged continue to grow laterally and upward. As the coral gradually grows upward and abruptly dies down through cycles of RSL rise and fall, the microatoll produces a distinctive concentric-ring surface morphology. In any given year, the highest elevation to which a coral grows is the highest level of growth (HLG); in contrast, the height above which corals die during a diedown marks the highest level of survival (HLS; Meltzner et al., 2010). The sensitivity of a coral microatoll to ELW

allows the coral to be a precise proxy of RSL changes. A cross-section slab of a microatoll provides a record of annual coral growth with precise HLS and HLG elevations with which to reconstruct a RSL curve.

Determining the cause of specific diedowns is an important first step in pinpointing drivers of RSL in a region. The spatial extent and magnitude of diedowns can help determine their causes. Where coeval diedowns are recorded across a broad region and have a similar magnitude, it suggests oceanographic or climatic influences. In contrast, Meltzner and Woodroffe (2015) note that diedowns with amplitudes of >10 cm and whose amplitude varies significantly over short distances (i. e., tens of kilometers) are most likely tectonically driven. With varying amplitudes of diedowns over a short distance, particularly in a known tectonically active area, diedowns can be used as paleoseismic markers to estimate coseismic slip, rupture extents, and the magnitude of past events (e.g. Philibosian et al., 2014, 2017). Although the diedowns themselves are triggered by the abrupt lowering of RSL, the relative elevations of successive diedowns in a microatoll reflect longer-term trends in RSL, which can result from regional to local processes and are due to secular sea-level changes, land-level changes, or a combination of both.

### 4. Methods

#### 4.1. Reconstructing relative sea level from coral microatoll cross sections

Cross sections of a representative living coral microatoll were extracted and used to reconstruct RSL at Almeida. We observed multiple living coral microatolls and documented their morphologies (Fig. 1B, also shown in Supporting Fig. S2), noting the number of concentric rings on each and their respective ring widths. We used a Leica TCRP1203 R300 total station, with an accuracy of 3 mm, to survey elevations of each ring, with a minimum of four points measured per ring. ALME-L14, with the greatest number of well-preserved concentric rings, was
anticipated to have the longest RSL record at the site and was selected for slabbing. Since the coral was living, age-dated samples were not required; the slab was extracted in November 2018 and annual growth bands could be counted backward.

We followed the methods outlined by Meltzner and Woodroffe (2015) to prepare the cross section for analysis. Slabs provide information on the internal structure and annual growth of the coral microatoll, allowing us to recover critical details of diedowns that can be obscured by erosion or overgrowth, out-of-sequence inward growth, on the coral surface. The ~10 cm thick slab was cut into several ~13 mm slices, with all slices scrutinized to assess which diedowns were recorded in each slice. Because different parts of the record were better preserved in different sections of each slice, two slices were x-rayed and interpreted. The x-rays of each slice were mosaicked together, and the annual banding and growth unconformities were traced digitally using a vector graphics editor.

Ideally, a coral microatoll used for a study of past RSL is found in its growth position, having never tilted, settled, or moved. At Almeida, the living microatoll with the best record, ALME-L14, was tilted; the cross section required a tilting correction to restore the coral to its original position. Coral microatolls form concentric rings of approximately uniform elevation, with some natural variability due to differential growth rates or differential wave splash on different sides of the coral (Meltzner and Woodroffe, 2015). This natural variability produces a range of elevations over which the outer living edge grows. In Sumatra, the range was ±8 cm (2σ) (Meltzner et al., 2010), and at one site in the Leizhou Peninsula, China, the range was ±8.5 cm (2σ) (Yu et al., 2009). We would expect to find some or all the concentric rings on the coral's upper surface to be systematically tilted in one direction for a microatoll that tilted during or after its growth, i.e., one side of the coral is at a higher elevation. This elevation difference is more apparent when the coral surface is surveyed as there is a wide elevation range along the same aged ring crest (Supplementary Fig. S7). While differential erosion of a coral may also leave apparent differences in elevation on the surface of a microatoll, this explanation can be excluded in cases where the surface is minimally eroded and each ring is well preserved.

Since we could not confirm whether ALME-L14 remained in its original position or was also transported when it was tilted, we determined that an additional adjustment was necessary to restore the coral to its likely original elevation. We assumed that, before tilting and any settling, ALME-L14 was at the same elevation as other nearby living microatolls, tens to hundreds of meters away. After correcting for the slab's tilt, we set the elevation of the highest HLG (the ring crest) just before the 2016 diedown to match the weighted mean of similar points on other untilted corals at the site. In Fig. 3, this is the zero elevation datum. Details for the tilt correction are provided in Supporting Text S1 and Fig. S3.

The indicative meaning of the proxy, essential for RSL reconstructions, comprises the indicative range (the vertical range over which the proxy naturally forms) and a reference water level (RWL), which is the central tendency of the indicator, often in the form a tidal datum (van de Plassche, 1986). In this study, we use the highest HLG (the ring crest) prior to the last diedown in a coral microatoll to estimate the microatolls' indicative meaning. The HLG points prior to each dieown correspond to the crests of the concentric rings of a coral, and can be readily surveyed on a population of microatolls. Additionally, the highest HLG point prior to a diedown has been noted to represent the closest approximation to RSL (Meltzner et al., 2010). We use the mean elevation of the ring crest prior to the 2016 diedown on untilted living coral microatolls at the site as the RWL in our reconstructed time series; the standard deviation is used for the indicative range. After correcting the slab elevations for tilting, the digitized cross sections were placed in the same reference frame as the field survey, and all elevations are presented relative to the RWL. Details on the indicative meaning are included in Supporting Text S2.

RSL reconstructions from microatolls provide precise age and

elevation data that can be used as sea-level index points (SLIPs) and marine limiting data, though the detailed RSL reconstruction from a coral microatoll is nuanced. A SLIP defines the position of RSL at a particular location and time, and must include information regarding the proxy's elevation, indicative meaning, and age (Shennan, 2015; van de Plassche, 1986). In contrast, marine limiting data provide only a minimum bound on the position of sea level at a particular time. Following Meltzner et al. (2010), Meltzner et al. (2012), and Philibosian et al. (2014, 2017), we treat the highest HLG point prior to each diedown as a SLIP, as it is the best estimator of RSL, and all other points in our time series as marine limiting data.

The calculation of RSL in any given year can be defined as:

$$RSL = E_c - RWL$$

where $E_c$ refers to the elevation of the coral in any given year, be it an HLG or HLS point, relative to the RWL, the mean elevation of the highest HLG prior to the most recent diedown at the site, in 2016. We consider three sources of elevation error. These include errors associated with the elevation of the past HLG or HLS point (± 10.0 cm, 2σ), uncertainty associated with the RWL (± 8.3 cm, 2σ), and any surveying error (± 2.0 cm, 2σ) which we added in quadrature. For age uncertainties, the annual banding provides precise age estimates but we include a conservative ±0.5 year (2σ) uncertainty to the timing of the index points. These errors are elaborated in additional detail in Supporting Text S3.

## 4.2. Modelling relative sea-level estimates and rates

We applied an Errors-in-Variables Integrated Gaussian Process (EIV-IGP) model from Cahill et al. (2015) to estimate RSL elevations and RSL rates using the SLIPs derived from the coral record. The EIV-IGP model uses an inverse method where a Gaussian process (GP) prior is placed on the RSL rate and then the GP is integrated to determine RSL estimates (Cahill et al., 2015). The errors-in-variable (EIV) statistical method accounts for age and vertical uncertainties when estimating RSL data and reconstructing the RSL curve (Cahill et al., 2015; Dey et al., 2000). This model does not include marine limiting data points, so only the SLIPs based on the highest HLG point prior to each diedown are used to model RSL rates in the coral.

## 4.3. Determining annual low water and lowest astronomical tide (LAT)

To compare the elevation of coral diedowns with tidal datums, to compare the timing of coral diedowns with low-water events, and more generally to validate the RSL record reconstructed from a coral microatoll, we use hourly tide-gauge observations from San Fernando, ~22 km south of the Balaoan site, as an independent record of RSL. To place the San Fernando observations in the same vertical reference frame as the coral microatolls, we compare the San Fernando data with those from a pressure-sensor tide gauge we deployed during fieldwork in 2017 and 2018, which we surveyed relative to coral elevations. To compensate for the gaps in the tide-gauge data, we attempt to reconstruct the missing data using a tide model, sea-level anomalies (SLA) from satellite altimetry, and a dynamic atmospheric correction (DAC). From this, we found the elevation and timing of the annual low water (ALW) of each year and we quantified the tidal datums for the site, including the lowest astronomical tide (LAT), which is the lowest predicted tide over the 18.61-year nodal tidal cycle. Details of these methods are in Supporting Text S4-S7.

# 5. Results

## 5.1. Coral microatolls in Almeida

A population of living Porites microatolls was documented at Almeida, Balaoan (Fig. 1B, Supporting Fig. S2). Many of these
microatolls had three or more visible concentric rings, including the outer living ring, and a center; the general surface morphology of these corals included a lower center and a higher outer living ring. The majority of the *Porites* microatolls were in a lagoon, roughly aligned behind a rampart. Despite this environmental set up, we observed no signs of ponding, usually identified by flat-topped microatolls due to the RSL stability in the pond. The outward increase in ring elevation of the microatolls confirms that these microatolls were responsive to RSL change and were not ponded behind a moat.

A few *Porites* microatolls (>1.5 m diameter) were found tilted in situ. Of these tilted microatolls, we surveyed and documented two. On each microatoll that had tilted, portions of the outer living rim were lower than the living rims of untilted living *Porites* microatolls at the site, indicating which side of the microatoll had dropped; however, on some tilted microatolls, including ALME-L14, which was carefully documented, a portion of the outer living rim was higher than the living rims of untilted living *Porites* microatolls at the site, indicating that the microatoll had pivoted in place, and the now-higher side of the microatoll had risen. The range of elevations for ring 1 on ALME-L14 was ±24.2 cm (2σ) as compared to the ±10.0 cm (2σ) across the untilted living corals (Supplementary Fig. S7). Notably, no post-tilting diedown was observed on the living rims of any tilted microatolls, suggesting that the polyps that rose had not yet been continuously exposed for a sufficient duration to kill them, but there were signs of bleaching on the crest of the tilted portion of ALME-L14, indicating that those coral polyps were experiencing heat stress on the days of our field work. This suggests that our site visit occurred soon after the tilting event, before a sufficiently low spring tide triggered any new diedowns on the portions of the living corals that had risen. Despite being tilted, the ALME-L14 microatoll (Fig. 1B, Supporting Fig. S2c-f) provided the most and best-

preserved concentric rings of any living microatoll at the site — a total of six rings including the living ring — so we chose it for slabbing.

## 5.2. Coral microatoll RSL reconstruction from ALME-L14

We use the two slices of the ALME-L14 slab and obtained a continuous RSL record from 1906 to 2018 CE (Fig. 2). A total of 17 diedowns were recorded in the slab: in late 1918, early 1925, early 1948, early 1958, early 1966, early 1973, early 1983, late 1987, early 1992, late 1992 or early 1993, early 1998, early 2004, mid-2005, mid-2006, early 2008, early 2010, and early 2016 (Table 1). The age uncertainty on these diedowns is conservatively ±0.5 yr (2σ) and, given these uncertainties, it is likely that diedowns in late 1992 in slice ALME-L14–1 and early 1993 in slice ALME-L14–2 resulted from a single diedown event. We categorize the diedown extents as small (≤3 mm) or large (>3 mm), and where a growth unconformity near the top of the coral does not involve the highest coral polyps living at the time, diedowns are determined to have a zero amplitude (Supporting Table S7). The coral growth rate varied across the slab, from 8 to 12 mm/yr, with a median growth rate of 10.5 mm/yr.

As noted earlier, ALME-14 had recently tilted when we found the coral, and when we attempted to correct for this tilting, we observed that the innermost concentric ring was tilted significantly differently than the outer rings, implicating an earlier tilting event in the coral's lifetime. A detailed analysis (Supporting Text S1 and Fig. S3) indicates the earlier tilting event occurred in 1948 ± 1 CE. For each tilting event, we restore the coral back to its original horizontality by resolving the dip or tilt of the relevant concentric rings until they are horizontal. For the more recent tilting event, we can then restore the coral (and each ring) to its original absolute elevation by comparing the ring elevations on the

<figure><img src="image_6.png" /><figcaption>**Fig. 2.** (Upper and lower left) Annotated cross sections of the two slices from the ALME-L14 slab (slices 1 and 2) from Balaoan. Annual banding and unconformities are traced where possible from the X-rays. The annual band ages were counted backward from 2018 CE, when we slabbed the living microatoll. (Right) Microlog of annual bands that are eroded in each slice such that evidence of a diedown might not be preserved, and the Ensemble Oceanic Niño Index of ENSO events, with El Niño (≥ 1) and La Nina (≤ -1) (Webb and Magi, 2022). Timing of diedowns in the coral record in dashed blue lines. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)</figcaption></figure>
<figure><img src="image_7.png" /><figcaption>**Fig. 3.** (A) Relative sea-level reconstruction from two slices of the ALME-L14 slab at Balaoan. Elevations are relative to the weighted mean of the ring crest prior to the 2016 diedown on untilted coral microatolls at Balaoan, assuming ALME-L14 tilted in place but was not transported and did not settle. The modeled RSL estimates are derived from SLIPs using the highest HLG prior to the diedown (orange squares). The modeled RSL rate represents the overall rate from the EIV-IGP model for the modeled time period. (B) Uniform vertical (± 13.2 cm) and chronological (± 0.5 yr) uncertainties (all at 2σ) are shown here at scale and apply to all RSL data on this graph.</figcaption></figure>

restored coral to corresponding ring elevations on nearby corals that were never tilted, and by uniformly shifting the restored coral up or down to minimize the misfit. For the earlier tilting event, we were unable to determine the original absolute elevation of the pre-1948 portion of the coral because no contemporaneous rings could be found on "untilted" corals at the site. Because of these complications, rates of RSL change can be estimated from this coral prior to 1948 CE, but the rates before and after 1948 must be estimated separately; RSL itself cannot be estimated from this coral prior to 1948 (Supporting Fig. S19-S20). From two slices of ALME-L14 we were able to reconstruct rates of RSL change at Almeida from 1906 to 1948 CE and from 1948 to 2018 CE, and we reconstructed absolute elevations of RSL from 1948 to 2018 CE.

Following the correction for tilting, SLIPs from HLG points prior to each diedown on ALME-L14 were used to estimate RSL positions and rates. A total of 15 SLIPs were used to determine the RSL change

between 1948 and 2018 CE. The vertical and chronological errors associated with each SLIP are determined to be ±13.2 cm (2σ) and ± 0.5 yr (2σ), respectively (Supporting Fig. S19). The EIV-IGP method provides a statistical framework to model RSL elevations and rates for each time step given the data available and their associated vertical and chronological errors (Supporting Fig. S19 and S20); from the time-varying rates we estimate the overall RSL rate for 1948–2018 CE as 1.3 ± 1.8 mm/yr (2σ). Prior to 1948, three diedowns are recorded in the slab. While we cannot know how far below ELW the coral was prior to the first diedown as unrecognized uplift may bias the coral elevation (Meltzner et al., 2010), fitting a curve through HLG points prior to each of the three pre-1948 diedowns nonetheless provides a maximum bound on the average rate of RSL rise for 1918–1934 CE, 15.1 ± 41.8 mm/yr (2σ) (Supporting Fig. S20A).
<figure><img src="image_8.png" /><figcaption>**Fig. 4.** (A) The San Fernando tide-gauge record overlying the reconstructed water levels. Elevations are adjusted to mean sea level at Balaoan. Also shown are the timing and elevations of diedowns from the ALME-L14 RSL reconstruction. The ALW from the reconstructed water levels and the tide gauge are plotted in (A) as circles. (B) Comparison of the ALW from the tide-gauge record and the reconstructed water levels, where possible.</figcaption></figure>

## 5.3. Tide gauge and reconstructed water levels

### 5.3.1. San Fernando tide-gauge record

The tide gauge at San Fernando has a record spanning from 1998 to the end of 2020, but has a completeness index of only 40.3 %, indicating that more than half of the data are not present (Supporting Fig. S8a). Focusing on the instrumental period from the beginning of 1995 to the end of 2020 from this record, the completeness index is only marginally better at 44.7 %. Noting the incompleteness of these data, we estimate an average rate of RSL of $-2.1 \pm 0.9$ mm/yr (Supporting Fig. S24) between 1995 and 2020 CE after correcting for any phase delays or periods with vertical offsets where possible (Supporting Text S7, Supporting Figs. S10-S18). We determined the ALW from the tide-gauge record but only 12 ALW could be robustly obtained for the 26-year period of 1995–2020 CE.

### 5.3.2. Reconstructed water levels of Almeida, Balaoan

The reconstructed water levels for Almeida provide a complete and hourly-resolution water-level record to replace the incomplete San Fernando tide-gauge data. We could reasonably reconstruct the tide-gauge record back to 1993 CE as satellite altimetry is needed to estimate SLA and DAC, but we limit our water-level reconstruction to 1995 onward due to a gap in the tide gauge between 1990 and 1995 that precludes validation of the reconstruction earlier than 1995. Comparing the tidal observations at San Fernando to the reconstructed water levels,

the mean absolute error is $7.4 \pm 11.1$ cm ($2\sigma$) (Supporting Text S7, Figs. S10–S18). The 12 ALW determined from the tide-gauge data were compared to the ALW estimated from the reconstructed water levels. The differences in the timing of the ALW between both datasets range from zero hours up to 55 days (Supporting Table S4). The timings of six ALW from the reconstruction are within four days or less of the timing of ALW from the tide-gauge observations (Fig. 4B). Five ALW determined from the reconstructed water levels are between 26 and 55 days off from the timing of ALW from the tide-gauge data (Fig. 4B). This 1 to 2 month offset in ALW is within the 3-month window of December through February when we expect lowest water levels to occur. When comparing the vertical misfits between the ALW from the reconstructed water level and the San Fernando tide gauge, the vertical difference is $6.8 \pm 14.7$ cm ($2\sigma$).

In reconstructing water levels and modelling tides for our site, we found some data inconsistencies in the San Fernando tide-gauge record that required adjustment. Tides follow a consistent periodicity that is not governed by climate phenomena, i.e. typhoons, or by land-level changes. A tide-gauge record, which records water level, is subject primarily to the tides and secondarily to other signals such as eustatic-sea level changes, ocean-atmosphere phenomena, and land-level changes. The periodicity of tides recorded by a tide gauge should remain consistent. When comparing the tide-gauge record to our reconstructed water levels, we found inconsistencies and errors in the tide-gauge record that required correction or removal. Around 80.3 % of the existing
<figure><img src="image_9.png" /></figure>

**Fig. 5.** (A) Simulated growth scenario of ALME-L14 following the reconstructed tides, showing periods when upward coral growth intersects low water levels. The solid and dashed sub-horizontal lines indicate the range of coral growth at minimum, median, and maximum rates of 8.0, 10.5, and 12.0 mm/yr, as the coral grows from the ALW of 1995. Diedowns from the ALME-L14 record are plotted as boxes with their ±2σ range (magenta boxes). Based on the water levels and coral growth, we denote where we expect and observe diedowns, and whether those diedowns were large or small in the slab. The corals were slabbed in 2018, so we did not observe the 2020 expected diedown. (B) The timing of El Niño (≥ 1) and La Niña (≤ −1) events based on the Ensemble Oceanic Niño Index (ONI) of Webb and Magi (2022). Dashed vertical grey lines cutting through (A) and (B) represent the timing of observed diedowns. (For interpretation of the references to colour in this figure legend, the reader is referred to the web version of this article.)

tidal observations are likely reliable, but for the other 19.7 %, these segments of the tide-gauge record either required temporal or vertical shifts or were removed. These are detailed in Supporting Text S7 and shown in Figs. S10–S18.

## 5.4. Relationship of living corals to the tides and annual low water

The highest elevation of HLG prior to the most recent diedown from the ALME-L14 slab was compared to the tidal datums estimated for the site using the tide model derived from the tide-gauge data. The elevation of HLG in 2015 generally lies between LAT and MLW (Supporting Fig. S7). The weighted mean of the HLG prior to the most recent diedown is 31.5 ± 10.0 cm (2σ) above LAT or 36.4 ± 10.0 cm (2σ) below MSL (Supporting Fig. S7). The living outer rim of eight *Porites* microatolls were also surveyed in 2018 to determine the elevation range of the living HLG; however, these survey points may not accurately represent RSL as the coral may have still been growing up to low water levels since its most recent diedown two years earlier.

The HLS following diedowns observed in the slab can be compared to the lowest water levels that coincided with those diedowns, where tidal records are available. Two diedowns coincide with suitably complete portions of the San Fernando tide-gauge record (Supporting Table S5). Using only these two diedowns and the tide-gauge record, we estimate HLS at 20.7 ± 2.0 cm (2σ) above observed ELW. Although this

relationship is based on very few data, the consistency (as indicated by the low standard deviation) suggests that a robust and precise relationship may exist between HLS and the elevation of the ELW that causes the diedown; more data would be needed to support or refute this argument.

We can potentially expand this comparison between HLS and ELW using the reconstructed water levels. In contrast to the two diedowns that coincide with tidal observations, with the reconstructed water level we can compare seven observed diedowns during this period with their modeled ALW (Fig. 5A). These HLS are 26.1 ± 11.7 cm (2σ) above the modeled ALW which coincides with a diedown (Supporting Table S6). Although the precision of this relationship appears to deteriorate with the inclusion of additional data, the increased scatter may arise from limitations of the water-level reconstruction (namely, limitations of the satellite-based SLA), and not from real variability in HLS as it relates to ELW. This inference is supported by the 6.8 ± 14.7 cm misfit (discussed in Section 5.3.2) between the ALW from the reconstructed water levels and the ALW observed in the San Fernando tide-gauge record. Additional observations are needed to resolve this question.

Using the continuous reconstructed water levels and the range of growth rates observed in the ALME-L14 slab, we can also predict the timing of diedowns expected in the coral record. In earlier work, Meltzner et al. (2017) did this for microatolls on the Sunda Shelf, assuming a reasonable range of coral growth rates and water levels that
<figure><img src="image_10.png" /><figcaption>**Fig. 6.** Absolute sea level, measured from satellites in space (A), and relative sea level, measured from corals on land (B) from 1995 to 2018 CE. (A) Sea-level anomalies for the nearest gridded point to Balaoan, with respect to a twenty-year mean (1993–2012). (B) The corresponding coral microatoll relative sea-level reconstructions with the overall EIV-IGP modeled rate. Differences between A and B, beyond uncertainties, are presumably due to vertical land motion, which is picked up in land-based observations, but not in observations from space.</figcaption></figure>

approximately followed a tide model. Here we build upon this strategy, starting in 1995 CE during the ALW, and we simulate the coral growth from this point onward using the reconstructed water levels (Fig. 5A). When the upward growth of the coral catches up to ELW, the coral dies down (indicated by the downwards vertical line) and after the low-water event ends the coral continues growing upwards. Following this method, diedowns are expected in 1998, 2002, 2003, 2004, 2005, and 2016 (Fig. 5A).

Comparing these predictions to the diedowns interpreted from the slab, there is reasonable agreement but also notable inconsistencies. Large (> 3 mm) and small (≤ 3 mm) observed diedowns (Supporting Table S7) are indicated by differently colored circles below the curve in Fig. 5A. Based on the reconstructed water levels, diedowns were both expected and observed in 1998, 2004, 2005, and 2016. In three instances (2006, 2008, and 2010) a diedown occurred that was not expected from the reconstructed water levels, with 2010 experiencing a large diedown, while 2006 and 2008 were small (Fig. 5A). In contrast, two expected diedowns are not observed in the slab: an early 2002 diedown clearly did not occur, but erosion of the 2003 growth band in both slices renders the existence of an early 2003 diedown unresolvable (Fig. 5A).

## 5.5. Vertical land motion at Almeida

We compare absolute sea level, measured from satellites in space, and relative sea level, measured from coral microatolls on land (Fig. 6);

details on the method can be found in Supporting Text S8. Trends estimated from satellite-based SLA (Fig. 6A) indicate sea-level rise at rates of 2.9 ± 0.5 mm/yr (2σ) between 1995 and 2018 CE. Over similar periods, the EIV-IGP modeled rates from the coral microatolls suggest RSL rise at higher rates of 3.0 ± 12.1 mm/yr (2σ) between 1995 and 2018 CE (Fig. 6B). Differencing the RSL and sea-level rates for the similar periods, we estimate VLM rates of -0.1 ± 12.1 mm/yr (2σ).

We cannot quantify VLM using the SLAs derived from satellite altimetry data and our reconstructed water levels, as reconstructed water levels are not independent of the SLA data. If we attempt to estimate VLM by comparing the satellite-based SLA data to the tide-gauge record, with its low completeness index, the VLM suggest uplift of 4.3 ± 0.8 mm/yr (2σ) between 1995 and 2021 CE (Supporting Fig. S24). This is primarily due to the low RSL rate derived from the San Fernando tide-gauge data at -2.1 ± 0.9 mm/yr between 1995 and 2020 CE.

# 6. Discussion

## 6.1. Correlation between coral diedowns and El Niño

El Niño is a climate phenomenon that occurs when the easterly trade winds weaken over the Pacific. As a result, the normally warmer ocean surface waters in the tropical western Pacific shift eastward, away from the Philippines (e.g. Bjerknes, 1969; Wyrtki, 1975). The weakening of the trade winds, and the replacement of buoyant warm ocean surface water off the Philippines by cooler, denser ocean surface water, leads to
a lowering of the sea surface in the equatorial western Pacific. A marked reduction of SSH during El Niño is well documented just east of the Philippines (e.g. Nerem et al., 1999), but the effect is less pronounced or more complex off the west coast of Luzon (Chang et al., 2013; Landerer et al., 2008; Pei et al., 2015). Nonetheless, lower sea levels tend to be observed at the San Fernando tide gauge during El Niño years (Fig. 5), leading us to anticipate a possible correlation between El Niño and coral diedowns.

Going back to 1918, the year of the earliest diedown on ALME-L14, the 12 strongest El Niño events on record occurred in 1918–19, 1930–31, 1940–41, 1957–58, 1965–66, 1972–73, 1982–83, 1987–88, 1991–92, 1997–98, 2009–10, and 2015–16 (Webb and Magi, 2022). Ten of those El Niños have a clear correspondence with a diedown on ALME-L14; the only strong El Niños without clear evidence of a diedown on the coral are 1930–31 and 1940–41 (Table 1). In the case of 1940–41, erosion of the upper parts of the 1940 and 1941 growth bands may have removed evidence of a diedown that occurred, so this case is ambiguous. Only one strong El Niño, in 1930–31, clearly lacks a corresponding diedown in ALME-L14.

Likewise, large diedowns, ranging in amplitude from 8 mm up to 89 mm, were recorded on ALME-L14 in late 1918, early 1925, early 1948, early 1958, early 1966, early 1973, early 1983, late 1987, early 1992, early 1998, early 2004, early 2010, and early 2016 (Fig. 2; Supporting Table S7). With the exception of the diedowns in early 1925, early 1948, and early 2004, all of these coincided with a strong or super El Niño (Supporting Fig. S21). The causes of the diedowns in early 1925 and early 2004 are unclear: the early 1925 diedown occurred during a La Niña event but preceded a moderate El Niño by less than one year, whereas the early 2004 diedown coincided with low water at the site, even if not during an El Niño. There are no storm tracks or signs of storm-induced tilting in 1925 and no records of significant earthquakes in the nearby region could be associated with uplift that could kill the upper parts of these corals. There are a few recorded minor coseismic events in 1924–1925 CE (details in Supporting Table S8) which could have resulted in some ground shaking and minor slumping in Balaoan (Southeast Asia Association of Seismology and Earthquake Engineering, 1985). In contrast, the early 1948 diedown appears to relate to a tilting of the coral, which will be discussed in Section 6.2. Additional smaller diedowns, with amplitudes less than 3 mm, were recorded on ALME-L14 in late 1992 or early 1993, mid-2005, mid-2006, and early 2008, but

<caption>**Table 1**</caption>

Diedown timing from the ALME-L14 record and the timings of “strong” to “super” El Niño phases based on the Ensemble Oceanic Niño Index (Webb and Magi, 2022). The El Niño type, Central Pacific or Eastern Pacific El Niño, is based on Freund et al. (2019).

<table><thead><tr><th rowspan="2">Diedown timing (year)</th><th colspan="2">El Niño events<fnref n="1" /></th><th rowspan="2">El Niño type</th></tr><tr><th>Start (year)</th><th>End (year)</th></tr></thead><tbody><tr><td>late 1918</td><td>1918</td><td>1919</td><td>Eastern</td></tr><tr><td>early 1925</td><td>—</td><td>—</td><td>—</td></tr><tr><td>early 1948</td><td>—</td><td>—</td><td>—</td></tr><tr><td>early 1958</td><td>1957</td><td>1958</td><td>Eastern to Central</td></tr><tr><td>early 1966</td><td>1965</td><td>1966</td><td>Central</td></tr><tr><td>early 1973</td><td>1972</td><td>1973</td><td>Eastern</td></tr><tr><td>early 1983</td><td>1982</td><td>1983</td><td>Eastern</td></tr><tr><td>late 1987</td><td>1987</td><td>1988</td><td>Central</td></tr><tr><td>early 1992</td><td>1991</td><td>1992</td><td>Central</td></tr><tr><td>late 1992</td><td>—</td><td>—</td><td>—</td></tr><tr><td>early 1993</td><td>—</td><td>—</td><td>—</td></tr><tr><td>early 1998</td><td>1997</td><td>1998</td><td>Eastern</td></tr><tr><td>early 2004</td><td>—</td><td>—</td><td>—</td></tr><tr><td>mid-2005</td><td>—</td><td>—</td><td>—</td></tr><tr><td>mid 2006</td><td>—</td><td>—</td><td>—</td></tr><tr><td>early 2008</td><td>—</td><td>—</td><td>—</td></tr><tr><td>early 2010</td><td>2009</td><td>2010</td><td>Central</td></tr><tr><td>early 2016</td><td>2015</td><td>2016</td><td>Eastern</td></tr></tbody></table>

these did not coincide with strong El Niños.

The influence on the timing and amplitude of diedowns by other factors such as solar insolation, temperature, and the duration of sub-aerial exposure are not well understood. High insolation, high temperatures, and a markedly decreased level of precipitation during El Niño in this region (e.g. Van Oldenborgh et al., 2021) may independently or collectively contribute to a diedown at times of fairly low water levels. Determining the incipience of diedowns would provide context to the specific drivers that influence the proxy elevation. Nonetheless, the close agreement between large diedowns and strong (or super) El Niño events demonstrates that coral microatoll diedowns are caused primarily by low water level, supporting their use as precise proxies for past RSL.

An extended period (multiple years, perhaps a decade or longer) of more frequent and/or stronger El Niño events can erroneously alias into local linear sea-level trends, biasing interannual to interdecadal estimates of secular sea-level change. This is due to correlations between El Niño and changes in SSH. In the South China Sea off the western Luzon coast, El Niño is found to correlate with lower SSH (Fang et al., 2006; Pei et al., 2015; Peng et al., 2013; Zhou et al., 2012). Annual mean sea level was found to lower by 5 cm during the 1997–1998 El Niño event in the eastern South China Sea (Fang et al., 2006). During El Niño events, SLA between 1993 and 2020 fell between 8 cm and 19 cm relative to mean sea level (Supporting Fig. S22). Zhang and Church (2012) used satellite altimetry to estimate 6 ± 4 mm/yr of apparent sea-level rise between 1993 and 2009 off western Luzon, largely a result of more frequent and stronger El Niño events between 1993 and 1998 than in subsequent years. Peng et al. (2013) had similar findings, but showed that sea-level reconstructions that assimilate tide-gauge data yield lower estimates of apparent sea-level rise for 1993–2009 near the Philippines, when compared to estimates strictly from satellite altimetry.

El Niño can be characterized by two models, the canonical Eastern Pacific (EP) El Niño and the Central Pacific (CP) El Niño. Diedowns occur during both of these modes. When categorizing the El Niño events correlated with diedowns, we find four diedowns occur during EP El Niño events, five during CP El Niño events, and one during a mixed event (an event that started as an EP and ended as a CP), as shown in Table 1 (Freund et al., 2019; Yu and Kim, 2013). The EP mode has been shown to have minimal influence on low-frequency variability in SSH in the SCS, whereas the CP type can have a more pronounced effect (Pei et al., 2015). Comparing the sea level-response in different El Niño models can be used in future work to understand which events could produce higher magnitude diedowns and RSL fall.

The overall strong correlation between El Niño and coral diedowns in western Luzon provides a framework for understanding what is recorded by coral microatolls in the region. The majority of the diedowns — especially those with large amplitudes of 8 mm or more and without evidence of an abrupt, persistent change in relative sea level — are likely to result from interannual fluctuations in sea level due to the El Niño–Southern Oscillation. Diedowns that seem to be associated with a permanent lowering of RSL, where subsequent diedowns are significantly and consistent lower on the coral, may result from tectonic uplift, but further analysis in these cases is generally warranted. Similarly, one or more diedowns followed by an extended period of unrestricted upward coral growth, with routine diedowns resuming only many years later and at a significantly higher level on the coral, may result from tectonic subsidence, but again further analysis is warranted. Coral microatolls respond to long-term sea-level changes and are limited by their annual growth rates (Yu et al., 2009) making it imperative that an appropriate reference elevation, e.g. the highest HLG prior to a diedown, is selected when constructing RSL time series. Long-term rates of change in the elevations of coral diedowns, averaged over many routine (El Niño) diedowns, offers insight into long-term rates of RSL change at the site, which may result from tectonics, GIA, localized subsidence due to compaction or groundwater withdrawal, or other processes.

<footnote n="1">¹ Based on El Niño events defined as “strong” or “super” (K ≥ 1.5) from Webb and Magi (2022).</footnote>
## 6.2. Coral tilting: typhoons or land movement?

The RSL reconstruction of ALME-L14 required correcting the tilt in the coral record, but what caused living corals at Balaoan to tilt to such a degree? A coral microatoll can be tilted as a result of slumping or settling in place on the substrate that the coral grew on. Proposed causes of slumping or settling include liquefaction or sediment compaction during strong ground shaking; coral microatolls on some sites in Simeulue and Nias Islands in Indonesia tilted during the 2004 Aceh-Andaman M<sub>w</sub> 9.2 and earlier earthquakes (Meltzner et al., 2015; Meltzner et al., 2012; Meltzner et al., 2010). In western Luzon, only three earthquakes above M<sub>w</sub>7.0 occurred between 1906 and 2018 CE, and none coincide with a diedown or tilting event in the ALME-L14 record (U.S. Geological Survey, 2017). Alternatively, corals can be lifted and moved, but this requires a significant force, particularly for a coral that is ~3 m in diameter. Balaoan has no recorded tsunamis or seiches in the 20th to early 21st century (Bautista et al., 2012), which rules out potential seismically induced wave action that could move the coral microatolls. Another plausible explanation is storm surge due to a tropical cyclone. Extreme storms and tropical cycles are known to push and transport large reef rocks and boulders (e.g., Goto et al., 2011; Kennedy et al., 2017; Scaffin, 1993). Similar observations are found across the South China Sea: there are documented emplaced coral reef boulders in Pasuquin, Philippines (Gong et al., 2022) and Hainan, China (Zhou et al., 2021, 2025), and emplaced boulders in Batangas, Philippines (Gallentes et al., 2021) and Penghu Islands, Taiwan (Yu et al., 2023). The wave energy generated by such storms can be destructive to reef environments and erode reef platforms (Done, 1992; Ryan et al., 2018).

We can reasonably attribute the most recent tilt of ALME-L14, in 2018 CE, to a known tropical cyclone. We conducted field work at Balaoan, including surveying and slapping of the microatolls, 24–28 November 2018. Importantly, during field work, we noted that ALME-L14 was tilted, and one side of the coral was exposed at low tide (Supporting Fig. S2). No untilted living microatolls at the site were exposed at these low tides, and the side that was exposed was bleaching but had not experienced a diedown. These observations suggest that the exposed side of the coral had risen when the coral tilted, but that the tilting was recent enough that the coral had not yet experienced a sufficiently low tide to cause a diedown. The tilting almost certainly had happened after the lowest tides of January 2018, which would have been lower than the low spring tide of 27 November 2018, which we witnessed during our site visit. Typhoon Yutu (locally known as Typhoon Rosita) made landfall in eastern Luzon on 29 October 2018, and weakened in intensity as it headed westward, crossing the Cordillera mountain range (Philippine Atmospheric, Geophysical and Astronomical Services Administration, 2020). Typhoon Rosita tracked directly over Balaoan, La Union, on 30 October (Philippine Atmospheric, Geophysical and Astronomical Services Administration, 2020) as a Category 1 storm on the Saffir–Simpson scale.

We argue that Typhoon Rosita offers a reasonable explanation for the tilting of ALME-L14 and can explain why no new diedowns had occurred since the tilting event. We assume ALME-L14 tilted and could have been transported during this process. A typhoon with sufficient energy can transport, tilt and, potentially, overturn microatolls. However, only tilted and no recently overturned microatolls were found at the site during our visit. The coral was stable after tilting, indicating ALME-L14 was on a stable substrate if it had moved. As we could not resolve whether ALME-L14 had moved or not, we assumed the mean elevation of ALME-L14 pre-tilting was consistent with other microatolls at the site.

A diedown observed in the coral slab at the beginning of the 1948 growth band is also attributed to a tilting event because this diedown separates an inner concentric ring, which is tilted in one direction, from the outer concentric rings, which are tilted consistently in a different direction. The diedown would have necessarily occurred after the tilting, on the next occasion when water was sufficiently low, likely during the spring tides of December or January following the tilt, as these tend

to be the months with the lowest tides (Fig. 5). Assuming the ±0.5 yr uncertainty in the growth banding, the tilting may have occurred in either 1947 or 1948 CE. Historical typhoon records suggest a few possible typhoons that could have tilted the ALME-L14 coral in those years. Two tropical cyclones made landfall in La Union in November 1947: Typhoons Dora and Flora (Supporting Fig. S23; Knapp et al., 2010, 2018). Both typhoons reached Category 3 on the Saffir Simpson scale (Knapp et al., 2018; Knapp et al., 2010). Alternatively, if the tilting occurred during the typhoon season of 1948, Typhoons Gertrude and Beverly tracked over the Almeida site in August and December 1948, respectively (Supporting Fig. S23; Knapp et al., 2010, 2018). These were weaker typhoons, with Gertrude reaching Category 2 and Beverly reaching Category 1. Presumably, at least one of these typhoons was strong enough to tilt the coral, which in 1948 CE was ~1 m in diameter and ~50 cm high in the center (Fig. 2, Supporting Fig. S23).

## 6.3. Feasibility of reconstructed water levels for sea-level studies

The reconstructed water levels provide a high temporal resolution dataset with which to compare the coral microatoll RSL record but must be used cautiously when applied to sea-level research. This approximation of RSL has large uncertainties when compared to the ALW of the San Fernando tide gauge (6.8 ± 14.7 cm; Supporting Table S4) that could not be resolved. The misfit may be due to differences in environmental setup: specifically, the San Fernando tide gauge is in Lingayen Bay (Fig. 1A), whose bathymetry and coastal runup may induce a local sea-level high, whereas the SLA record used for the reconstructed water levels is based on a gridded point from the open ocean. These differences are significant, and unless improved, limit the utility of the reconstructed water-level record for direct comparison to the coral microatoll RSL record. The model could be optimized by careful selection of SLA data. It is noted that satellite altimetry data, either gridded or retracked, with the highest correlation to the tide-gauge record are observed at a point 300 km from the tide gauge, and are not as sensitive to distance as expected (e.g. (Oelsmann et al., 2021; Peng et al., 2024); on the flip side, our chosen SLA point (the nearest grid point to the tide gauge) may not be particularly representative of the San Fernando tide gauge. Where possible, a long-term tide gauge located at the study site of interest would provide the closest approximation and independent record of sea-level changes.

## 6.4. 20th-century rates of sea level and VLM in western Luzon

Long-term sea-level changes have been modeled to a limited extent in the regions surrounding western Luzon and the South China Sea. Regional sea-level reconstructions in the South China Sea use mixed models of satellite altimetry and tide gauges, extending the tide-gauge record (Church et al., 2004; Peng et al., 2013). These reconstructions suggest sea-level rates estimated at ~1.5 mm/yr between 1950 and 2000 CE (Church et al., 2004) and 1.5–2.5 mm/yr between 1950 and 2009 CE (Peng et al., 2013). The advent of satellite altimetry provides an independent measure to quantify sea-level change since 1993. We modeled SLA rates between 1995 and 2018 at 2.9 ± 0.5 mm/yr (Fig. 6A). Rietbroek et al. (2016) report a faster sea-level rise between 2002 and 2014 CE for the South China Sea at 7.6 ± 3.17 mm/yr. However, their rate spans a shorter period and is not constrained to sea-level changes along western Luzon. Modeled sea-level rates from mixed tide gauge–satellite altimetry datasets show lower rates: the “virtual” tide gauges modeled by Cazenave et al. (2022) suggest sea-level rates offshore western Luzon of roughly 2–3.5 mm/yr between 2002 and 2019 CE; separately, Peng et al. (2013) report sea-level rise of ~5 mm/yr from 1993 to 2009 CE using a mixed tide gauge–satellite altimetry dataset for the South China Sea.

The RSL rates modeled from the coral microatoll cannot be directly compared to sea-level rates derived purely from SLA or SSH changes, as these do not capture any local VLM present in the RSL record. As a
tectonically active region that has not experienced a major earthquake in recent history, interseismic deformation is a key driver of VLM at our site. We attempt to quantify VLM using the RSL record from coral microatolls but recognize that improvements may aid the calculation of a more precise VLM rate. The SLA rates were chosen as the closest gridded point to the site; however, this requires tide-gauge and satellite-altimeter data to be of the same temporal resolution. Our VLM rates based on a comparison between SLA-derived rates of absolute sea-level change (Fig. 6A) and RSL rates estimated from the corresponding coral microatoll RSL record (Fig. 6B) assimilate rates from the EIV-IGP model and are therefore sensitive to the length of time over which the rate is modeled and the number of SLIPs during that period (Cahill et al., 2015). For the period seen in Fig. 6A and B of 1995–2018, the VLM rate is quantified at −0.1 ± 12.1 mm/yr (2σ). The high RSL uncertainty in the RSL record between 1995 and 2018 in Fig. 6B is due to the limited number of SLIPs in the short period of time modeled. Due to the sparsity of data in the San Fernando tide gauge, the modeled RSL rate was not stable, but it is shown in Fig. S21A.

Based on the surface morphologies of the living coral microatolls, we can infer that the contributions of VLM in the region are relatively minor. The increase in the concentric ring elevation over time acts as an independent measure that RSL was rising during the coral's lifetime. The modeled rates from SLIPs derived from the slab record also show a positive RSL trend with some uncertainties that could reflect RSL fall. The SSH from satellite altimetry indicate a rising sea level. Therefore, the rates of VLM should be less than the SSH trend, which could represent either a low uplift rate or low long-term subsidence. We present estimates of VLM in this section, understanding that the approximate rates do indicate tectonic deformation may occur in the region, but also that reducing the uncertainties in the modeled RSL rates would aid in constraining the VLM rates in the region.

## 7. Conclusions

This study presents a 20th-century relative sea-level reconstruction from coral microatolls in western Luzon and finds long-term RSL rise of 1.3 ± 1.8 mm/yr between 1948 and 2016 CE. During the satellite observational period, from 1995 to 2018 CE, the rate of relative sea-level from the coral microatoll record is 3.0 ± 12.1 mm/yr. We attempt to quantify rates of vertical land motion using the nearest gridded sea-level anomaly point and the relative sea-level record from the coral microatolls and estimate interseismic rates of −0.1 ± 12.1 mm/yr for the 1995–2018 CE period. The rate of vertical land motion indicates potentially low interseismic rates in the region. Tropical cyclones play a major role in tilting coral microatolls and explain the tilting we observed in the ALME-L14 coral microatoll record. In the absence of earthquakes, diedowns in the region are controlled by extreme low water during El Niño events. Indeed, the close agreement between diedowns and El Niño at our site demonstrates the sensitivity of coral microatolls to low water level and supports their use as precise proxies for past RSL. The diedown timing is also controlled by the growth rate of the coral microatoll up to its theoretical HLS. The reference water level, the weighted mean of the ring crest prior to the most recent diedown, is found to lie 31.5 ± 10.0 (2σ) cm above the lowest astronomical tides in an 18.61-year nodal cycle between 2000 and 2018 CE. Comparing the two instances where diedowns overlap with the tide-gauge record, we estimate HLS to be 20.7 ± 2.0 cm above ELW. Alternatively, using the reconstructed water levels, we determine HLS to be 26.1 ± 11.7 cm (2σ) above ELW using seven diedowns. We present a method of reconstructing water levels by combining a local tide model, sea-level anomalies, and a dynamic atmospheric correction, and validating this reconstructed water level to a local water level survey and a nearby tide gauge. The reconstructed water levels can be used for determining modern rates of relative sea-level change.

## CRediT authorship contribution statement

**Gina M. Sarkawi:** Writing – review & editing, Writing – original draft, Visualization, Validation, Project administration, Methodology, Investigation, Formal analysis, Data curation, Conceptualization. **Aron J. Meltzner:** Writing – review & editing, Supervision, Project administration, Methodology, Investigation, Funding acquisition, Conceptualization. **Dongju Peng:** Visualization, Resources, Methodology, Investigation. **Joanne T.Y. Lim:** Writing – review & editing, Investigation, Formal analysis. **Xinnan Li:** Writing – review & editing, Investigation, Formal analysis. **Rohan Gautam:** Writing – review & editing, Validation, Software, Investigation, Formal analysis. **Anandh Gopal:** Writing – review & editing, Investigation. **Andrew Mitchell:** Writing – review & editing, Investigation. **Loraine Faye Sarmiento:** Writing – review & editing, Investigation. **Jennifer Weil-Accardo:** Investigation. **Kathrine V. Maxwell:** Writing – review & editing, Investigation. **Junki Komori:** Writing – review & editing, Investigation. **Noelynna T. Ramos:** Writing – review & editing, Project administration.

## Declaration of competing interest

The authors declare that they have no known competing financial interests or personal relationships that could have appeared to influence the work reported in this paper.

## Acknowledgements

We thank the barangays of Paraoir and Almeida, Balaoan for allowing us to conduct fieldwork on their land, in particular, thank you to R. Agnes, W. Madiam, R. Morales, G. Pacleb. We thank the Department of Agriculture Bureau of Fisheries and Aquatic Resources for allowing us to slab the coral microatolls. We are grateful to our cutting team including R. Cortez, L. Tiston, G. Quina, G. Tiston, L. W. Orsolina, D. Prayudi, and Amar for helping collect data. We thank N. Tiongson, P. Cortez, D. Leoterio, for fieldwork driving and assistance. We are grateful to A. G. Dee at the Teresa Marble Corporation for allowing us to slice our coral microatoll slab at their facility. We are appreciative of J. T. Catabui at the University Health Service at the University of the Philippines Diliman for allowing us to x-ray the coral microatoll at their facility. We thank P. Flores and the Physical Oceanography Department at the National Mapping and Resource Information Authority of the Philippines for access to tide-gauge data. We thank N. Cahill for assistance with the Errors-in-Variables Integrated Gaussian Prior models. We are appreciative to E. Hill, J. Freymueller, F. Tan, D. Lallemant, and M. Rabonza for useful suggestions. We thank two anonymous reviewers for their constructive feedback, which greatly improved this manuscript. This work was primarily funded from the National Research Foundation (NRF) Singapore under its Singapore NRF Fellowship scheme (Award NRF-NRFF11-2019-0008 to AJM). This research is also supported by Singapore Ministry of Education (MOE) Academic Research Fund (Award MOE2019-T3-1-004) and the National Research Foundation (NRF) Singapore and the Singapore Ministry of Education (MOE) under the Research Centres of Excellence initiative. This is Earth Observatory of Singapore contribution no. 637.

## Appendix A. Supplementary data

Supplementary data to this article can be found online at https://doi.org/10.1016/j.margeo.2025.107565.

## Data availability

Data is available through the DR-NTU (Data) research data repository at https://doi.org/10.21979/N9/32SKMK. These include field data (field notes, field photos, slab photos, raw x-ray images, annotated slab mosaics, survey files), time series, coral tilt correction, tide-gauge
data and the associated files and scripts for determining tides and
reconstructing water level.

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