Sea level changes forced ice breakouts in the Last Glacial cycle: new results from coral terraces

Quaternary Science Reviews 21 (2002) 1229–1240

Sea level changes forced ice breakouts in the Last Glacial cycle: new results from coral terraces John Chappell* Research School of Earth Sciences, Australian National University, Australian Geological Survey Organ, Mills Road, Canberra, ACT 0200, Australia Received 21 December 2000; accepted 31 August 2001

Abstract Sea level cycles recorded by coral terraces at Huon Peninsula (HP) in Papua New Guinea match rapid climate changes that occurred between 30,000 and 65,000 years ago, seen in Greenland ice cores and marine sediments. Each cycle of 6000–7000 years ended with a sea level rise of 10–15 m lasting 1000–2000 years, following a longer period of falling sea level. Precise dating shows that each rise corresponds to a ‘‘Heinrich’’ episode of ice-rafted detritus in north Atlantic, signalling massive ice outbreak from north America. Sea level may be the trigger that forced near-synchronous breakouts of ice from north America and eastern Greenland. The HP sea level changes also correspond to similar cycles in benthic oxygen isotopes reported from the north Atlantic, which are too large to be explained entirely in terms of ice volumes. Jointly, the sea level and isotope records suggest that the north Atlantic deep ocean cooled as sea level fell and warmed as sea level rose, in each 6000-year cycle. r 2002 Published by Elsevier Science Ltd.

1. Introduction It is well established that large, millenial-scale climatic shifts occurred repeatedly in many parts of the northern hemisphere during the Last Glacial period, particularly from about 12,000–75,000 years ago (12–75 ka). These are recorded by oxygen isotope fluctuations known as Dansgaard–Oeschger (DO) cycles in Greenland ice cores (Grootes et al., 1993; Stuiver and Grootes, 2000), by microfauna and isotopes in sediment cores from various marine basins including the north Atlantic (Bond et al., 1993; van Kreveld et al., 2000), the Mediterranean (Cacho et al., 1999) and the Santa Barbara Basin (Kennett et al., 2000), and by particle size variations in north China loess and Sea of Japan sediments (Porter and An, 1995; Tada and Irino, 1999). As seen in the Greenland ice cores, a typical DO cycle has an average period of B1460 years, with a cold phase of B600 years that terminates with an abrupt switch to a warmer phase. Isotopically, the amplitude of a typical DO cycle is about 50–75% of the full glacial–interglacial range (Stuiver and Grootes, 2000). Ice-rafted detritus (IRD) in marine sediments show that ice breakouts from Green*Tel.: +61-262-295-111; fax: +61-262-257-247. E-mail address: [email protected] (J. Chappell). 0277-3791/02/$ - see front matter r 2002 Published by Elsevier Science Ltd. PII: S 0 2 7 7 - 3 7 9 1 ( 0 1 ) 0 0 1 4 1 - X

land precede abrupt DO warmings (van Kreveld et al., 2000). In the north Atlantic marine sediments, DO cycles have been grouped in bundles known as Bond cycles that terminate with IRD horizons known as Heinrich events, from massive ice outbreaks from Labrador (Bond et al., 1993). The simultaneous occurrence of rapid changes in regions far from the north Atlantic ice fields indicates wide-ranging links in the climate system. Possible mechanisms reviewed by van Kreveld et al. (2000) range from luni-solar forcing to the effects of massive iceberg flotillas that interrupt the north Atlantic thermohaline circulation and lead to changes in the other oceans. Linkages and feedbacks within the climate system range from methane pulses from the oceans (Kennett et al., 2000) to aridity-driven fluctuation of atmospheric dust (Broecker, 2000). Whatever the causes, ice sheets are involved but possible behaviours range from collapse and surge of unstable sheets through slower cycles of ice growth and decay. The amplitude, rate and timing of the resulting sea level changes depend on the ice breakout mechanism. This paper shows that between 30 and 65 ka the Bond cycle bundles of DO cycles were locked into sea level changes that are recorded in raised coral reefs at Huon Peninsula (HP), Papua New Guinea. The flight of

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downstepping reefs at HP includes broad terraces with ancient lagoons and barriers that formed during major sea level rises at terminations of major glacial cycles, which are separated by groups of smaller fringing-reef terraces produced by sea level fluctuations of various scales within glacial periods. Earlier studies from HP identified sea level changes at orbital periods (20– 100 ka), supported by U-series dates of comparatively low precision (Bloom et al., 1974; Chappell, 1974) but did not examine the millenial-scale record. Sea level determinations from the HP terraces have the advantage

of not being affected by temperature changes, which complicate the interpretation of sea level from marine oxygen isotope records. This paper derives sea levels from precise topographic and stratigraphic data supported by published high-precision U-series ages.

2. Coral terraces and sea level changes Fig. 1 shows two transects of Late Pleistocene reefs at HP: Kanzarua (KANZ) and Bobongara (BOBO). These

Fig. 1. Sections of coral terraces II–IIIau at BOBO (above) and KANZ (below), based on detailed surveys of topography (Chappell et al., 1996b) and bio- and litho-facies and stratigraphy (Chappell, 1974; Pandolfi and Chappell, 1994); K1, BU10, etc. show locations of dated corals (Chappell et al., 1996a) listed in Table 1. Key to facies: 1, reef platform; 2, reef crest and buttress; 3, reef slope; 4, lagoon limestone. Basal surface of reef deposits is interpolated between observations of underlying tuffs, conglomerates and bioclastic sediments of the Bobongara Beds (key 5), seen in ravine exposures beneath reefs II and IIIa and in surface exposures between IIIau and IV. Blank areas=reef deposits not seen. Terrace numbering follows Chappell et al. (1996a); arrows indicate shore angle at rear of terraces. Small unamed terraces and notches represent stillstands between tectonic uplift events.

J. Chappell / Quaternary Science Reviews 21 (2002) 1229–1240

are based on previous reports of the stratigraphy (Chappell, 1974; Pandolfi and Chappell, 1994), detailed topography (Chappell et al., 1996b) and TIMS U-series age measurements (Chappell et al., 1996a). The sections extend from Late-Glacial Reef I to Late Pleistocene Reef IIIau, dated at B60 ka. Reef I is a raised barrier and lagoon complex, up to 400 m wide, whereas reefs IIb–IIIau are relatively narrow and are capped by terraces of platform and beach deposits over coral limestone comprised mainly of shallow-water reef platform and reef crest deposits, that usually overlie forereef deposits (Chappell, 1974; Pandolfi and Chappell, 1994). Small unnamed terraces between the named terraces in Fig. 1 represent episodes of erosion and cliff retreat that occurred between uplift events, similar to erosion benches and wave-cut notches in the raised Holocene reef. Detailed dating of the Holocene benches shows that uplift at HP is dominated by metre-scale events with an average recurrence interval of about 1000 years (Ota et al., 1993; Chappell et al., 1996b). To illustrate the interpretation of sea levels from raised reefs, Fig. 2 shows simulated effects on terrace morphology, and reef facies of sea level changes shown as curves SA and SB : Both curves have highstands at 37 and 44 ka but SB has a smaller rise at 37 ka and a much faster rise at 44 ka. Sections SA and SB simulate the corresponding reefs and terraces (produced by the computer model described below, with tectonic uplift of 3 m/ka). Thicknesses of shallow-water reef crest and platform deposits (dark-shaded zones, Fig. 2) reflect amplitudes of relative sea level rise; terrace slopes and widths reflect both rates and amplitudes of rises. In section SA ; for example, the 44 ka terrace is the horizontal top of a shallow-water keep-up reef that kept pace with the relatively slow rise to the 44 ka highstand in curve SA : By contrast, the 37 ka terrace slopes seaward because its reef lags behind rising sea level except at the shoreline, so that the outer part of the terrace is underlain by shallowing-upwards deposits of a catch-up reef. As well as the difference in slope, Fig. 2 shows that the keep-up terrace has the highstand shoreline at its seaward margin, whereas the highstand shoreline is at the landward margin of a catch-up terrace. In section SB ; both rises were rapid and led to catch-up reefs with sloping terraces, and the 37 ka terrace now is only half as wide, because the rise to the highstand at 37 ka in SB is half that in SA : From Fig. 2, the thickness DZ of the reef crest and platform deposits beneath a given terrace reflects a sea level rise DS: DS ¼ DZ þ U Dt;

ð1Þ

where U is the uplift rate and Dt is the time interval from low to highstand. Dt cannot be determined without

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Fig. 2. Terrace morphologies and reef facies respond to sea level changes: sections SA and SB correspond to sea level curves SA and SB ; respectively (sections are produced by the computer model described in the text, with tectonic uplift=3 m/ka). Shallow-water reef crest and platform deposits underlie the terraces (dark-shaded zones); sloping lines represent submarine reef surfaces at sea level highstands and lowstands. Terrace slopes and widths reflect rates and amplitudes of sea level change (see the text).

accurate dating but is constrained by reef type: keep-up reefðhorizontal terraceÞ; DtXDZ=G0 ;

ð2aÞ

catch-upðsloping terraceÞ; DtoDZ=G0 :

ð2bÞ

Parameter G0 is the potential growth rate of the reef crest, which is known for Holocene reefs at HP from dated drill-core and is constrained by total reef thicknesses, as described below. Rewriting 2(a) and (b) as equations by multiplying a parameter Cc (with Cc ¼ 1 for keep-up; Cc o1 for catchup) and substituting (1) gives the rate of sea level rise: DS=Dt ¼ G0 =Cc þ U:

ð3Þ

In practice, estimates of DS based on field sections are likely to be minimum values, because the full vertical extent (DZ) of crest/platform deposits occurs at buried transgressive surfaces and usually is not at exposed terrace fronts. Expressions (1)–(3) are used below to

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determine sea level changes relative to each reef at BOBO and KANZ. Dated corals within the reefs are used to reduce these relative sea levels to differences from present sea level at HP by subtracting tectonic uplift: the results are referred to as ‘‘HP sea levels’’. Thus, the HP sea level St represented by coral of age t at height H above sea level is St ¼ H þ D  Ut

with U ¼ ðHd  Sd Þ=td ;

minimum estimates of HP sea levels (St ) and field-based estimates of DZ for each terrace. Uplift rates used to calculate St differ slightly from previous values (these were U(BOBO)=3.2–3.4 and U(KANZ)=2.7– 2.9 m/ka, based on values of Sd ¼ 572 m and td ¼ 12273 ka (Chappell et al., 1996a). By analogy with the raised Holocene reef, where the crest formed at or soon after the end of the Late-Glacial sea level rise (Chappell et al., 1996b), the age of the Last Interglacial reef crest should not be much younger than the end of the Penultimate deglacial rise, dated at 128 ka with Sd ¼ 025 m (Stirling et al., 1998). With Hd ðBOBOÞ ¼ 405 m and Hd ðKANZÞ ¼ 340 m, and td ¼ 1262128 ka with Sd ¼ 025 m, gives UðBOBOÞ ¼ 3:1323:21 and UðKANZÞ ¼ 2:6222:7 m/ka. Table 1 also summarises the relationship of terrace nomenclature in Fig. 1 to previous HP naming schemes.

ð4Þ

where U is the uplift rate at the sample site and D is the water depth in which the coral grew (sign: positive downwards). U increases from northwest to southeast along the HP coast (Bloom et al., 1974; Chappell, 1974) but at any given transect the Holocene uplift rate and the rate since the Last Interglacial are very similar (Ota et al., 1993). Setting aside metre-scale uplift events, U is assumed to be constant at a given terrace transect and is calculated from the local height Hd of the reef crest formed in the Last Interglacial, when sea level Sd was similar to or a few metres higher than today (Chappell et al., 1996a). As well as support from the similarity of Holocene and Last Interglacial rates, this assumption is favoured by the finding, below, that sea level results derived from BOBO and KANZ are very similar, despite that their uplift rates are different. If uplift rates had varied through time, the assumption of uniform uplift would usually (although not necessarily) lead to different sea level reconstructions from different sections. Table 1 lists previously reported heights, ages and estimated water-depth ranges for corals from sample points shown in Fig. 1, together with maximum and

3. HP sea levels from KANZ and BOBO sections Sea level based on dated corals from KANZ and BOBO are graphed in Fig. 3, together with sea level changes based on DZ values for each reef, listed in Table 1. DS=Dt for each sea level rise is estimated using Eq. (3) with G0 as 10 m/ka, as found for the Late-Glacial reef at HP based on U-series dates from a 52-m drillhole (Edwards et al., 1993). Cc was taken as 0.9 because Cc must be o1 for catch-up reefs, and all except perhaps IIa at BOBO are the catch-up type. Sea level highstands in Fig. 3 are estimated from catch-up terrace heights at

Table 1 Age, height and sea level controls from BOBO and KANZ reefs II–IV Fig. 1a

COBb

CSb

Height H (m)

Age t (ka)

Error 2s (ka)

Depth D (m)

HP sea levelc Min

Max

30 30 49 26 28

33.4 33.0 37.8 34.8 42.1

0.6 0.5 0.3 0.3 0.03

2–5 2–5 0–2 2–5 1–3

78 77 74 67 86

68 67 67 60 79

IVa

BU21 BU24 BU10 KU14 KU9 F KU10 K34 K9 K4 K3 K1

49 56 78 86 96 105

43.9 44.5 54.6 51.2 51.8 61.4

0.7 0.7 0.7 0.8 0.8 0.6

5–15 2–5 2–10 2–5 1–3 0–2

67 64 70 53 46 64

50 54 54 42 35 54

F Vb

F FRT1

183

72.8

2.2

0–5

54

37

IIb

Iib

II

IIa

Iia

II

IIIc IIIb

IIIc IIIb

IIIb IIIa

IIIal IIIam

IIIa.1 IIIa.2

IVb

IIIau

IIIa.3

F IV

IIIa.4 IV

a

Sample codea

DZ (m)d

Dt (ka)d

4 (B)

o0.4

8 (B) 6 (K)

>0.8 o0.6

(B) (K) (B) (B) (K) (B) (K)

o1.0 o0.8 o0.4 o0.6 o0.8 o1.8 o1.4

8 (B)

o0.8

10 8 4 6 8 18 14

Fig. 1 terrace names, sample codes (abbreviated codes), heights and ages from Chappell et al. (1996a). COB=terrace names from Chappell et al. (1996b) for BOBO and KANZ. These are consistent with Fig. 1 but the numbering used by Chappell and Shackleton (1986) is offset, as shown under CS in column 3. c Smin ¼ H þ Dmin Umax ðt þ 2sÞ; Smax ¼ H þ Dmax  Umin ðt  2sÞ: See the text for U-values. d DZ estimates from KANZ (K) and BOBO (B) sections; Dt derived from DZ by using G0 ¼ 10 m/my (see the text). b

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Fig. 3. (Top) Sea level based on dated samples (points with error bars) and sea level rises (heavy lines) based on shallow-water reef thicknesses, from Table 1; light lines are preliminary interpolations between data points. Points and horizontal error bars are plotted at the bottom of vertical error bars, to emphasise that the sea level curve should not lie below these points but can lie above them, even beyond the vertical bars because water depth of coral growth is poorly constrained. Curve does not pass through samples KU14 and KU9 for reasons given in the text. (Below) Reef sequence at BOBO simulated by reef model driven by the sea level curve, above. Heavy shading indicates simulated shallow-water platform and reef crest deposits; sloping lines show seabed surfaces at highstands and lowstands. The observed BOBO profile is shown (laterally offset) for comparison.

their landward margins (cf. Fig. 2), and each highstand is assigned the age of the oldest dated sample from its corresponding reef. The figure represents the simplest construction that is consistent with the terrace stratigraphy of Fig. 1 and the dates in Table 1. Thus, to complete the curve, highstands are joined directly to their succeeding lowstands. By doing this, two samples do not lie on the sea level curve: (i) the line from the 44 ka highstand to the 39 ka lowstand passes above sample KU9 (42.1 ka) and (ii) the curve bypasses sample KU14 (34.8 ka) which is thought to be from reef IIa but may be from an eroded remnant of IIb. Fig. 3 does not show an oscillation for terrace IIIal, which is undated (dated sample K9 from below IIIal is older than the samples from IIIam and probably represents a lower part of reef IIIam). Fig. 3 is regarded as a preliminary sea level curve, owing to uncertainties in DZ; which arise because the best exposures occur at terrace fronts but greatest thickness of shallow-water reef tends to occur beneath

the central and landward sides of the terraces, and to uncertainties stated as inequalities in formulae (2). The exact relationship of most samples to the highstands is not known but the stratigraphy in Fig. 2 implies that corals at or below terrace margins will be somewhat younger than the related highstands, whereas samples from gullies and ancient sea cliffs may be either older or younger than their highstands. Furthermore, the value of G0 ¼ 10 m/ka comes from the Late-Glacial barrier reef at HP (Edwards et al., 1993), but lower rates have been found elsewhere and this value may not be applicable for the Late Pleistocene fringing reefs at BOBO and KANZ. To test the Fig. 3 curve, it was fed into a computer model that simulates coral reef growth under varying sea level with tectonic uplift. Simulations are tested against the KANZ and BOBO sequences. In this model as in previous models (Chappell, 1980; Bosscher and Schlager, 1992), the potential rate of vertical accretion of coral limestone is assumed to be constant (G0 ) from

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close to the sea surface down to a critical depth (Dc ) that represents the limit of photo-compensation, commonly about 15 m (Bosscher and Schlager, 1992). Below Dc ; coral and sediment accretion is assumed to decrease linearly with increasing depth and is zero below a cut-off depth Dz ; which is set at 100 m on the basis of presentday carbonate sediments offshore at HP and on the stratigraphy of bioclastic limestone in deeper parts of the Pleistocene HP reefs (Chappell, 1974; Pandolfi and Chappell, 1994). Simulations with Dz ¼ 50 m did not differ significantly, in terms of terrace and shallow reef structures. The inititial substrate was taken as a seaward-dipping uniform slope, which approximates the surface overlain by the offlapping reef sequences at HP (Fig. 1). Uplift (U) is taken as uniform, and metrescale uplift events are neglected. At each time step, sea level is obtained from the input curve; uplift is incremented and the new reef profile is computed explicitly from the prior profile, commencing at the deep-water end. Graphic output shows the simulated terrace topography, reef sequence and reef profiles for successive highstands and lowstands. Simulations based on sea levels in Fig. 3 show both points of agreement and points of disagreement with observed terrace sequences. Fig. 3 shows the simulation for BOBO with G0 ¼ 10 m/ka (as was used to construct Fig. 3) and U ¼ 3:2 m/ka. This model reproduces terrace IIa and IIIb at about their observed heights and widths, but simulated terraces IIb and IIIam are too narrow; IIIal is absent, and IIIau is too high and does not slope seawards. The KANZ simulation showed similar discrepancies. Results suggest that the estimated sea level rises for misfit terraces are too small (terrace width tends to increase with the amplitude of sea level rise, Fig. 2). Furthermore, in this simulation the thickness of the reef stack above the prior surface is much greater than observed. Better simulations were sought by varying the sea level curve. First, G0 was reduced so that reef-stack thickness matched observations: G0 ¼ 4 m/ka gave the best results and values up to 6 m/ka were acceptable. Alternatives were examined with diferent values for lowstand sea levels, rates of rise (DS=Dt) and highstand sea levels. For each rise, the highstand–lowstand difference DS was varied from zero to twice that shown in Fig. 3. Rise rates were varied from B7 (cf. Eq. (3) with G0 ¼ 4; Cc ¼ 0:9 and U(KANZ)=2.6 m/ka) to >30 m/ka (i.e., Cc o0:2). Highstand ages controlled by dated samples (terraces IIb, IIa, IIIb, IIIam, IIIau) were not varied. Transitions from highstands to their succeeding lowstands were varied by introducing intermediate inflection points. To assess a sea level experiment, the fit between simulated and observed profiles was evaluated in terms of terrace heights, widths and slopes, thicknesses of shallow-water reef, and age-height results for dated

samples, with a ‘‘% goodness of fit’’ index (f ): X X At =mÞ%; f ¼ 20ð ðAh þ Aw þ As þ Az Þ=n þ

ð5Þ

where Ah ; Aw ; As ; Az and At are nominal (1; 0) ‘‘agreement’’ scores for terrace heights, widths, slopes, shallow reef thicknesses (DZ), and dated sample ages, respectively, n is the number of terraces and m is the number of dates in consideration. Agreement is accepted (score 1) when observed and simulated terrace heights lie within 72 m, widths are within 720%, slopes are within 721 and are >01 for catch-up terraces, and shallow reef thickness is 0–5 m greater than observed DZ: For a dated sample, At ¼ 1 if its age-height error bars overlap the modelled sea level curve. The score for the simulation in Fig. 3 is f ¼ 63%; by comparison, the BOBO simulation shown in Fig. 4, which was generated with the sea level curve in Fig. 4, has an f -score of 96%. By experimentally varying the sea level curve as outlined above, a restricted envelope of curves was found that generate high-scoring simulations (f > 90%) for BOBO and KANZ together. The best-fit envelope (Fig. 4) passes through all dated samples except KU14, which could not be reconciled with the need for a rise prior to reef IIb, nor was it possible for simulations from a single curve to match both the slope of catch-up reef IIa at KANZ and the horizontal platform of IIa at BOBO. High f -scores of acceptable simulations show that these are small details. The best-fit sea level envelope (Fig. 4) differs from the initial curve (Fig. 3) in that it has a lowstand between IIIau and IIIam and a small oscillation between IIIam and IIIb (this generates terrace IIIal), and the lowstand curves prior to IIa and IIb are not sawtooth-shaped. Timing of highstands remains tied to the dated samples. In terms of amplitudes of the rising phases, the best-fit curve is conservative because larger sea level rises at greater rise rates could generate the observed terrace widths. However, rise rate is constrained by terrace slope. With reef growth G0 of 4–6 m/ka, the best match to observed terrace slopes was with sea level rise rates of 8–10 m/ka (Cc ¼ 0:8 in Eq. (4)), but as the rate increases above 12 m/ka, simulated terraces rapidly became too steep. To summarise, Table 2 lists the timing and minimum estimate of the amplitude for each sea level rise, for reefs IIb–IIIau.

4. HP sea levels and global climatic events According to Fig. 4, the HP reef sequence IIb–IIIau represents six sea level oscillations with amplitudes ranging from B9 (IIb) to >20 m (IIIau), indicating that substantial fluctuations of ice volume occurred between 30 and 65 ka (for comparison, melting the entire Greenland ice sheet today would raise sea level by about 5 m). Sea level changes vary around the world

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Fig. 4. (Top) Typical ‘‘acceptable’’ simulation of BOBO section compared with observed profile. (Centre) Envelope of HP sea level curves that give high-scoring simulations of both BOBO and KANZ (lower boundary of the envelope gives the BOBO simulation shown above). (Below) Highresolution GISP-2 oxygen isotope record on layer-counted time scale (from Stuiver and Grootes, 2000), with selected interstadials identified by numbers. Small shaded bars represent Heinrich events H4, H5 and H5.2, with timing tied to GISP-2 as given by van Kreveld et al. (2000).

because of global isostasy and a change relative to Pacific coasts such as HP will be less than the iceequivalent value (Yokoyama et al., 2000). Differences depend on the sea level history but for the relatively rapid oscillations seen in Fig. 4, the amplitudes will be close to ice-equivalent values. Comparison of HP sea levels with other records reveals relationships between ice volume and millenialscale climatic changes. Fig. 4c shows the layer-counted Greenland GISP-2 ice core oxygen isotope record, which arguably has the best chronology amongst published high-resolution palaeoclimate records for the period of interest (Stuiver and Grootes, 2000; van Kreveld et al., 2000). Fig. 4 shows that sea level highstands at 38, 44.5 and 52 ka coincide with DO interstadials 8, 12 and 14, within dating errors. Ice-

rafted detritus pulses represented by Heinrich events 4, 5 and 5.2 and by IRD from east Greenland, tied to the GISP-2 chronology (van Kreveld et al., 2000), lie within the periods of sea level rise prior to these highstands. The highstand shown at 60 ka, which is the climax of the largest sea level rise in Fig. 4, must represent the termination of marine isotope stage 4 (MIS4) and may correlate with interstadial 17 but is discussed later. The reef IIb highstand at 33 ka apparently corresponds to interstadial 5 or 6, or interstadial 7 if sample KU14 came from a remnant of reef IIb and not from IIa. Of all these HP sea level cycles, only reef IIIal has no obvious equivalent in GISP-2. Table 2 summarises the correlations. Links between the ice sheets, sea level and Bond cycles are to be expected (Bond et al., 1993; Broecker, 2000).

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Table 2 Timing of key sea level and abrupt warming events HP reefa Highstand age (ka) Lowstand age (ka) Highstand sl. (m) Minimum sl. rise DS (m) Dd18 OS ¼ kS DSb GISP-2 interstadialc DO warming (ka) Heinrich number IRD, start age (ka) IRD, finish age (ka) N. Atlantic benthic d18O (core MD95-2042d) Highstand (ka) Lowstand (ka) Highstand d18O% Lowstand d18O% Dd18O% Max. residuale (%) Max. warmingf (1C) Pacific benthic d18O (core V19-30g) Highstand d18O% Lowstand (%) Pacific Dd18O% Max. residual (%) Max. warming (1C) South Atlantic IRD Peak ageh

IIb

IIa

33 34 72 9 0.07 DO5 32.3 F 32.6 32.3

? F F F F F F

4.48 4.67 0.19 0.12 0.5 SA2 30.4

IIIb

38 40 71 9 0.07 DO8 38.4 H4 39.4 38.4

44.5 46 56 14 0.11 DO12 45.4 H5 45.8 45.4

IIIam 52 54 46 16 0.13 DO14 52.2 H5.2 53.5 51.5

IIIau 60 62 50 26 0.21 DO17 58.5 H6? F F

38.4 40.1 4.28 4.63 0.35 0.28 1.2

45.3 47.6 4.18 4.70 0.52 0.41 1.8

51.7 55.0 4.30 4.57 0.27 0.14 0.6

57.7 61.1 4.05 4.75 0.7 0.49 2.1

4.41 4.66 0.25 0.18 0.8

4.51 4.61 0.10 0.01 0.0

4.3 4.58 0.28 0.15 0.6

4.2 4.66 0.46 0.25 1.1

SA3 36.6

SA4 43.8

SA5 51–55

d18O

a

Sea level rises from Fig. 4b. For a sea level change DS; the equivalent marine isotope change Dd18 OS ¼ kS DS (see the text). c DO warmings from GISP-2 (Stuiver and Grootes, 2000); Heinrich events and combined Greenland–Labrador IRD phases (van Kreveld et al., 2000). d North Atlantic benthic data from Shackleton et al. (2000). Raw data shown in Fig. 5; tabled d18O values are 3-point means. e Max. residual=Dd18 O2kS DS (from row 5, above). f Max. warming=kT (max. residual) (see the text). g Benthic d18O from east Pacific core V19-30 with correlations from Chappell and Shackleton (1986) renumbered as in Table 1. Maximum deepwater (DW) warming=d18O residual. b

What is new in the HP results are the direct estimates of amplitudes and rates of ice volume changes and their timing. The rates of sea level rise (ice volume decrease), constrained by terrace widths and slopes and by the rate of shallow-water reef growth G0 ; are no faster than the Late-Glacial rise (when rates were up to B15–20 m/ka, Edwards et al., 1993) and are much less rapid than the abrupt DO warming events. Together, the durations and rates of these rises indicate that each ice retreat generally was progressive and was not a single catastrophic surge although a series of surges within each rise cannot be ruled out. Concurrent IRD deposition (Table 2) suggests semi-continuous ice breakout through much of each rising phase. Detailed evidence from marine cores shows that abrupt DO warming occurred when IRD deposition ceased (van Kreveld et al., 2000), which apparently

happened at or shortly before termination of sea level rise, according to Fig. 4 and Table 2. Between 35 and 55 ka, ice breakouts with B10–15 m sea level rise appear to occur only with DO events at Bond cycle terminations (setting aside the relatively small oscillation of reef IIIal). To date, no sea level rises have been recognised at HP for most of the other DO cycles in this period. However, because uplift occurs in metre-scale events at HP (Ota et al., 1993; Chappell et al., 1996b), it would be difficult to identify sea level cycles of o2–3 m. The pattern of ice breakout with IRD deposition followed by abrupt warming occurs in all the other DO cycles (van Kreveld et al., 2000), so sea level presumably did oscillate in each cycle but by less than the HP detection limit of 2–3 m. This difference between Bond and DO cycles must reflect the relative sizes of

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their dominant ice breakout sources: marine sediments indicate that eastern Greenland is the major source in DO cycles (van Kreveld et al., 2000), which could not feed a 10 m sea level cycle, whereas Bond cycles involve the much larger Laurentide ice sheet (Bond et al., 1993). Interestingly, numerical modelling suggests that offpolar ice sheets of Laurentide dimensions would wax and wan at Bond cycle periods of about 6 ka without external forcing (Ghil, 1988). The joint Bond/DO terminations, marked by IRD from Greenland and the Laurentide sheets, show that ice breakout from both sources was nearly simultaneous (with Greenland apparently leading by a few hundred years, van Kreveld et al., 2000). This seems very unlikely for ice sheets of such different sizes, unless breakout was initiated by a common trigger. Atmospheric forcing is unlikely to be the trigger because SST and Greenland temperature rise at the end of each breakout, not at the beginning. However, a sea level rise could act as a trigger when both sheets are potentially unstable: a rise caused by breakout of one could initiate breakout of the other. Several factors affect the potential breakout under rising sea level, including relationships between ice-sheet base, underlying topography and the grounding line, which will vary with sea level and total ice mass and may be the basis of the fluctuating amplitudes of DO cycles seen in Fig. 4. Furthermore, the same mechanism might affect the margins of the Antarctic ice sheet. Extensive IRD pulses in southern Atlantic sediment cores are interpreted by Kanfoush et al. (2000a, b) as the result of major instability of the ice sheet margins, and not a stable response to climatic forcing. Calibrated radiocarbon ages of South Atlantic IRD events SA3 (36.6 ka), SA4 (43.8 ka) and SA5 (B51 ka) reported by Kanfoush et al. (2000a) lie close to sea level highstands represented by HP reefs IIa, IIIb and IIIam but appear to be 1–2 ka younger than H4, H5 and H5.2 in the north Atlantic (Table 2). However, new radiocarbon and U-series age data reported by Beck et al. (2001) suggest that the calibration corrections used by Kanfoush et al. (2000a) could be too small by B1 ka. If so, ice sheet instabilities in the north and south may be linked and sea level rise could be the common cause. The results presented here lead to a further point concerning sea levels and chronology. If we accept that Bond cycle terminations immediately follow Heinrich events that are coeval with sea level rises of 10+ m, then GISP-2 ages can be used when sea level events are imprecisely dated. This reasoning suggests that HP highstand IIIau=DO interstadial 17=58.5 ka, rather than the 60 ka age assigned from a single U-series date (Table 2). This age-shift requires somewhat higher sea levels for reef IIIau, and raises the highstand to 46 m at 58.5 ka and the upper limit of the prior lowstand to 72 m at 60.5 ka. Simulations of BOBO and KANZ

with these values are very similar to those based on the sea levels shown in Fig. 4 (f ¼ 94%).

5. HP sea levels and oxygen isotopes Correlation of an earlier version of HP sea levels and the benthic oxygen isotope record in east Pacific core V19-30 led Chappell and Shackleton (1986) to conclude that the deep ocean temperature was B1.51C cooler during the Last Glacial cycle than in the present or the Last Interglacial. In principle, variations of deep ocean temperatures during Bond cycles can be estimated in the same way. Huon Peninsula sea level cycles between 35 and 65 ka appear to match cycles in high-resolution benthic isotopic data reported by Shackleton et al. (2000) from north Atlantic core MD95-2042, which is age-locked to Greenland ice cores by DO events revealed in its planktonic isotope record. As shown in Fig. 5, the two records are very similar except that MD95-2042 has no obvious equivalent to the 33 ka (reef IIb) sea level peak and the brief isotope spike at B41 ka has no counterpart in the HP sea levels. Setting these aside, high HP sea levels for reefs IIa, IIIb, IIIam and IIIau correspond to low isotope values in MD95-2042, which appear as peaks on the inverted scale in Fig. 5, at 38.4, 45.3, 51.7 and 57.7 ka (adopting the 58.5 ka age proposed above for IIIau improves the age-correlation of this highstand and the 57.7 ka isotope peak). Low sea levels correspond with high isotope values at 40.1, 47.6, 55.0 and 61.1 ka. Notably, the benthic isotope record does not show the abrupt warmings seen in DO cycles but shows more gradual rises, similar to the sea level curve. MD95-2042 ages and isotopic values for peaks and troughs are listed in Table 2, which also gives isotopic values and peak-to-peak correlations of HP sea levels with benthic d18O in east Pacific core V19-30. The SPECMAP chronology of V19-30 is not precise and the correlations for this core in Table 2 follow Chappell and Shackleton (1986), with terraces renumbered according to the COB scheme in Table 1. According to the data in Table 2, each Bond cycle in Table 2 is marked by changes of oxygen isotopes and sea level. Possible effects of deep ocean temperatures can be assessed by using the following relationships: Dd18 O ¼ Dd18 OS þDd18 OT þ e; Dd18 OS ¼ kS DS; 18

ð6aÞ

Dd18 OT ¼ kT DT; 18

18

ð6bÞ

where Dd O ¼ d OðmaxÞ  d OðminÞ for a given cycle in a marine core, subscripts S and T refer to sea level and temperature, respectively, and e includes the effect of any changes on large-scale salinity gradients in the oceans. The temperature coefficient kT is 0.23%/1C (Shackleton and Opdyke, 1973), while coefficient kS depends on the isotopic composition of ice melted

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Fig. 5. Best-fit HP sea levels from Fig. 4, compared with benthic d18O record reported by Shackleton et al. (2000) for east Atlantic core MD95-2042. Heavy lines on the isotope record indicate isotope trends that apparently correspond to sea level rises.

(deposited) during a given sea level rise (fall). For the post-glacial sea level rise, DS is B125 m (Yokoyama et al., 2000) and Dd18O(sea level)=170.1% (Shackleton, 2000), giving a value of k ¼ 0:008%/m On this basis, maximum values of lowstand–highstand changes of deep water temperatures (DT) were estimated for each cycle, for both MD95-2042 and V19-30 using DS values derived from the HP reef sequence. Table 2 gives the results. In calculating DT; e is assumed to be zero because no changes in large-scale salinity gradients have been identified (Labeyrie et al., 1987). The deep sea DT estimates in Table 2 for rising sea level episodes IIa, IIIb and IIIam range to 0.81C for Pacific core V19-30 and to 1.81C for Atlantic core MD95-2042. For the relatively large sea level rise represented by reef IIIau (>26 m), Table 2 shows DT as 1.11C for the Pacific and 2.11C for the Atlantic. These values are quite large, compared with previous estimates for the full interglacial–glacial temperature shift in the deep ocean, which is considered to be B1.5–2.51C in the east Pacific and 3–41C in the equatorial Atlantic (Chappell and Shackleton, 1986; Labeyrie et al., 1987). The DT estimates would be smaller if the sea level rises were larger than the minimum estimates listed in Table 2, which is possible because recent analysis of

marine and atmospheric isotope records by Shackleton (2000) suggests that the sea level component of highstand–lowstand isotope variation of Dd18OS may have been B50% greater than the HP estimates in Table 2. These results could be reconciled if the value for kT was larger than the figure of 0.008%/m, used above, which also would reduce the estimates for DT: However, this would imply that the ice involved in these sea level cycles was isotopically lighter than the average for the glacialage ice sheets as a whole. The issue cannot be resolved here. To conclude, the chronologic and isotopic data in Table 2 strongly suggest that sea level highstands in the north Atlantic are preceded by deep ocean warming and lowstands are preceded by deep ocean cooling. The pace of these cooling and warming cycles matches that of the sea level (ice volume) changes, rather than the abrupt switches seen in DO cycles. Correlation with the benthic isotope record in core V19-30 indicates similar though smaller temperature changes in Pacific deep water but the amplitude and phase of these are uncertain, because the V19-30 record lacks precise age control. Indeed, if the sea level and isotope peaks were phase-shifted, DT estimates for V19-30 would be larger. Details found in Shackleton (2000), who shows temperature-dependent

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isotope fluctuations of up to 0.5% between 65 and 30 ka in east Pacific benthic records, suggest that this may have been so.

6. Conclusions A set of coral terraces dating between 30 to 65 ka at HP, Papua New Guinea, represent a series of sea level changes within the latter half of the Last Glacial cycle. Earlier studies reported topographic data, stratigraphic sections and precise U-series ages but did not analyse sea level changes on the basis of all available information. This paper approached the problem by the inverse method, using a computer model of reef growth driven by sea level. A restricted envelope of sea level curves was found that yield simulations that closely match the observed sequences at HP. The best-fit sea level envelope shows cycles with peaks at 33, 38, 44.5, 52 and 58–60 ka. The timing is fixed by precise U-series dates. Each cycle of 6–7000 years ended with a sea level rise of 10–15 m lasting at 1–2000 years, following a longer period of falling sea level. Except for the short cycle terminating at 33 ka, each sea level rise corresponds to a Heinrich episode marked by IRD in north Atlantic, signalling massive ice outbreak from north America, and the sea level highstands thus coincide with terminations of Bond cycles. Sea level rise may be the trigger that forced synchronous breakouts of ice from north America and eastern Greenland, and may also have triggered ice breakouts from Antarctica, although reported radiocarbon ages of the relevant IRD deposits in the South Atlantic are 1–2 ka younger than the north Atlantic events. Finally, the sea level cycles terminating at 38, 44.5, 52 and 58–60 ka coincide with similar cycles in benthic oxygen isotope reported by Shackleton et al. (2000) for north Atlantic core MD95-2042. The isotopic variations are larger than that can be accounted for in terms of sea level and suggest that temperature in the Atlantic deep ocean varied by 1–21C, in phase with the 6–7000 year sea level cycles.

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