Postglacial flooding history of Mayotte Lagoon (Comoro Archipelago, southwest Indian Ocean)

Marine Geology 194 (2003) 181^196 www.elsevier.com/locate/margeo

Postglacial £ooding history of Mayotte Lagoon (Comoro Archipelago, southwest Indian Ocean) J. Zinke a; , J.J.G. Reijmer a , B.A. Thomassin b , W.-Chr. Dullo a , P.M. Grootes c , H. Erlenkeuser c a

b

GEOMAR, Research Center for Marine Geosciences, Wischhofstrasse 1^3, Geb. 4, D-24148 Kiel, Germany Centre d’Oce¤anologie de Marseille (CNRS UA 41), Station marine d’Endoume, rue de la Batterie-des-Lions, 13007 Marseille, France c Leibniz-Laboratory for Radiometric Dating and Stable Isotope Research, Christian-Albrechts-University of Kiel, Max-Eyth-Strasse 11^13, D-24118 Kiel, Germany Received 26 October 2001; received in revised form 13 November 2002; accepted 13 December 2002

Abstract Four cores from the fringing reefs and five sediment cores from Mayotte Lagoon, Comoro Archipelago, southwest Indian Ocean, which reached the Pleistocene/Holocene boundary, form the database of this study. They offer the opportunity to reexamine and complete the postglacial sea-level curve, especially for the time interval between 11.6 to 8 kyr cal BP. Between 11.6 kyr cal BP until present the history of sea-level rise showed the following steps: (1) by about 19 mm/yr between 11.6 and 9.6 kyr cal BP, (2) by 9 mm/yr between 9.6 and 8 kyr cal BP, (3) by 3 mm/yr between 8 and 7 kyr cal BP, and (4) by 0.9 mm/yr after 7 kyr cal BP until stabilisation at present level at 2.5 kyr cal BP. In addition, a decline in the rates of sea-level rise or even a stillstand is observed between 13 to 11.6 kyr cal BP. The flooding of the lagoon of Mayotte was controlled by the depth of the reefal passages, which were cut by rivers and/or due to erosion during the time of emergence since the last interglacial. The differences in the shape of the sea-level curve from Mayotte compared to those from other sites located far from the former glaciated regions are related to: (1) the small size of the island, (2) the rapid downward movement of this small volcanic island with the oceanic plate into the mantle due to hydro^isostatic compensation after addition of meltwater, and (3) the location between large continents. A 2003 Elsevier Science B.V. All rights reserved. Keywords: Holocene; Younger Dryas; sea-level curve; fringing reef cores; sediment cores; southwest Indian Ocean

1. Introduction Mayotte is the largest barrier-reef-lagoon complex within the southwestern Indian Ocean. The lagoonal area measures approximately 1500 km2

* Corresponding author. E-mail address: [email protected] (J. Zinke).

with a maximum water depth of 80 m (Fig. 1). The length of the barrier reef is about 140 km and may reach a width of 2 km. The sediment distribution within Mayotte Lagoon and the growth pattern of the di¡erent types of reefs is mainly determined by the sea-level rise during the deglaciation from oxygen isotope stage 2 to stage 1. Deep passages were cut through the Pleistocene barrier reef by rivers or were formed

0025-3227 / 03 / $ ^ see front matter A 2003 Elsevier Science B.V. All rights reserved. doi:10.1016/S0025-3227(02)00705-3

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by karst processes since the time of emergence, which lasted from about 60 to 80 kyr cal BP (Camoin et al., 1997). Renewed £ooding of the lagoon occurred at ¢rst through these passages. The £ooding history of this conically shaped volcanic island with a pronounced older reef rim is controlled by the depth of these passages and the overall morphology of the pre-Holocene substrate. Mayotte Island thus forms an excellent example to study such a complex £ooding history. This study completes the sea-level history for a speci¢c time interval (11.6^8 kyr cal BP) through analysis of several cores from the lagoon and fringing reefs of Mayotte. We focus on the deglacial period between 11.6 kyr cal BP and the present, in which the meltwater pulse 1B occurred (Fairbanks, 1989). We discuss and compare our results with other Holocene sea-level studies (Atlantic: Fairbanks, 1989; Bard et al., 1990; Chappell and Polach, 1991; Indian Ocean: Eisenhauer et al., 1993; Camoin et al., 1997; Montaggioni and Faure, 1997; Yokoyama et al., 2000, 2001 ; Paci¢c: Pirazzoli et al., 1988; Blanchon and Shaw, 1995; Bard et al., 1996; Grossman et al., 1998; Hanebuth et al., 2000).

2. Previous studies The sea-level curve for Mayotte Island is based on the analysis of coral samples collected from the foreslopes using a submersible and by cores drilled into the eastern Pamandzi barrier reef (Fig. 2 ; Thomassin et al., 1993; Colonna, 1994; Colonna et al., 1996; Camoin et al., 1997; Dullo et al., 1998). The sea-level curve was constructed on the basis of 234 U/230 Th-datings and the estimation of palaeo-water depths during reef growth using biological assemblages (Montaggioni and Faure, 1997 ; Table 1). The average analytical uncertainty (1c) of these measurements is K 500 years for deglaciations and K 300 years for Holocene ages. Tectonic subsidence was estimated on the basis of the age of the karsti¢ed Pleistocene reef top occurring at a depth of 20 m below present sealevel (mbpsl) (Camoin et al., 1997). If the Pleistocene reef top corresponds to isotope stage 5a,

which would be a maximum assumption, than subsidence would range from 20 to 25 cm/ka resulting in subsidence of Mayotte of about 4^5 m during the last 20 kyr (Camoin et al., 1997). The early stages of sea-level rise, between 18.2 and 13.6 kyr cal BP, were moderate with rates of about 2 mm/yr. From 13.6 K 0.7 kyr cal BP to about 10 kyr cal BP, sea level rose rapidly reaching rates of about 25 mm/yr, which was related to meltwater pulse 1A (Dullo et al., 1998). A second rapid pulse was detected between 10 and 7.5 kyr cal BP with a rate of about 6 mm/yr. This event was related to meltwater pulse 1B (Dullo et al., 1998). During this time, vertical reef accretion was the fastest for the Holocene period with reef growth rates of 8 mm/yr, building up from 21.5 to 7 mbpsl (Camoin et al., 1997). Sea level rose by about 1.1 mm/yr between 7.5 and 2.5 kyr cal BP and stabilised around 2.5 kyr cal BP. Reef accretion slowed down to rates of about 1 mm/yr (Camoin et al., 1997). An even higher sea-level stand within isotope stage 1.1, as known from Paci¢c equatorial islands, the eastern Indian Ocean and Madagascar (Peltier, 1991; Pirazzoli et al., 1988; Pirazzoli and Montaggioni, 1988; Pirazzoli, 1991; Eisenhauer et al., 1993; Camoin et al., 1997; Grossman et al., 1998), is not recorded on Mayotte and was related to the small size and intraplate volcanic origin of the island (Camoin et al., 1997). The same holds for sea-level curves derived from the islands of La Re¤union and Mauritius situated in the southwestern Indian Ocean (Fig. 2 ; Montaggioni, 1979; Montaggioni and Faure, 1997; Camoin et al., 1997).

3. Materials and methods Twenty-seven conventional or AMS-radiocarbon dates on organic samples (mangrove root fragments, organic matter), bivalves, molluscs and in situ corals were produced at the LeibnizLabor in Kiel (Germany) and in the Laboratoire de Ge¤ologie du Quaternaire in Marseille (France) (Table 1). The U/Th-dates on in situ corals from the barrier reef core PMI-7 and the foreslopes were produced at the Bureau de Recherches Ge¤ologiques et Minie'res in Orle¤ans (France).

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183

M A DA G A SC A R

M O ZA M B IQ U E

C H A

N N EL

AFRICA

MAYOTTE

Passe M'Zamboro

Longoni lagoon 187

207 40

200

189

Passe Choizil

fringing reef cores C1-C8

186

20

40

201

Grand

190

89011

191

barrier reef

250

Grand*Terre

Grand Passe de l'Ouest

206 198

500

Boueni lagoon

20

205

*

60

Boueni Bay

Petite Terre

89027 89026

Passe Sada

500

Ajangoua lagoon

250

204

89028

Passe Boueni

PMI-7, barrier reef core

192

Passe Longogori

203 500

202 Passes de Saziley

193 89022

20

Passe aux Bateaux

barrier reef

197

194

Dapani lagoon

89011 = core number 20

= bathymetry

187-207 = submersible dive sites Fig. 1. Locality map of Mayotte Island situated within the northern area of the Mozambique Channel. The core locations (rectangles and asterisks) within the studied lagoonal areas are indicated by individual numbers. The submersible dive positions at the foreslopes of the barrier reef are indicated by small numbers.

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Fig. 2. Comparison between the reconstructed sea-level curves from the southwestern Indian Ocean (after Camoin et al., 1997) and the predicted sea-level curve for Re¤union Island based on model ICE 3-G (after Peltier, 1991).

Carefully screened and cleaned carbonate samples were selected for AMS dating and processed according to the methods described by Nadeau et al. (1997, 1998). The carbonate samples for conventional dating comprised coral fragments and mollusc shells of snails and oysters. Allochthonous sediments and secondary carbonates were carefully removed through soaking in water, hard ultrasonic agitation, high pressure water lancing, and mechanical brushing. Heavily encrusted parts were cut o¡, snail shells dissected alongside for cleaning, and secondary calcite coatings such as in worm tubes removed with milling cutters. About 25 K 3 g of the sample material was hydrolysed by HCl under 4.5-grade nitrogen to

produce CO2 . The organic samples were cleaned and humic acids were extracted (Nadeau et al., 1997). Extracted humic acids were also measured to test contamination by younger material. The age di¡erence between the humic acids and the organic material was less than 2c of the mean and, consequently, is not signi¢cant. AMS-14 C ages were measured with a 3-MV High-Voltage Engineering Europa Tandetron 4130 AMS system to a precision (counting statistics and machine statistics) of 0.3% for modern samples. This resulted in typical 14 C age uncertainties (1c) of 25 to 70 yr (Nadeau et al., 1997, 1998). Conventional 14 C dating was done on CO2 by low-level proportional counting.

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Table 1 List of radiocarbon dates obtained for the fringing reef cores and sediment cores from the lagoon of Mayotte Sample number

Depth in core (m)

Material

14 C-age (yr BP)

MAY.Kl.89011, water depth: 41.6 m, 12‡42P17QS, 45‡08P07QE LGQ622c 2.02^2.10 coral 8890 K 160 MAY.KL.89022, water depth: 26 m, 12‡58P80QS, 45‡09P08QE KIA5903 1.84 gastropod 1350 K 35 KIA5904 1.94 bivalve 2090 K 35 KI4386.001 2.30 bivalve 3880 K 35 KIA5905 2.60 coral 5970 K 50 KIA11560 2.80 organic 8740 K 45 MAY.KL.89026, water depth: 56 m 1.70^1.80 bivalve 5700 K 970 LGQ598c LGQ599c 3.55^3.62 bivalve 10190 K 190 LGQ600c 3.65^3.70 organic 9650 K 190 LGQ601c 4.45^4.49 organic 9860 K 210 5.20^5.50 organic 10270 K 410 LGQ602c MAY.KL.89027, water depth: 56 m KIA11558 3.80 organic 10000 K 40 KIA11559 4.95 organic 10070 K 45 MAY.KL.89028, water depth: 37 m, 12‡54P00QS, 45‡06P18QE KIA5204 0.50 organic 1870 K 30 KIA5205 1.13 organic 6460 K 30 KI4331.001 2.04 oyster 8750 K 70 KIA8739 2.73 coral 9020 K 100 KI4332.001 2.75 oyster 9150 K 95 KIA5206 2.97 organic 9420 K 40 Fringing reef core Longoni C1, water depth: 3.3 m KI4382,011 4.0^5.0 coral 1390 K 40 KI4282,021 9.0 coral 2550 K 35 Fringing reef core Longoni C4, water depth: 0.2 m KI4383,011 2.50 coral 2870 K 40 KI4383,021 10.0 coral 7510 K 45 Fringing reef core Longoni C5, water depth: 3.75 m KI4384,011 4.50 coral 2760 K 25 KI4384,021 8.0 coral 6780 K 55 Fringing reef core Longoni C8, water depth: 0.32 m KI4385,011 2.15 coral 5940 K 45 KI4385,021 7.30 coral 6470 K 45

Cal agea (yr BP)

13

Cb

Cal age (yr BP 2, c range)

Delta

9450 K 470

9923^8979

n.d.

900 K 30 1655 K 55 3850 K 40 6360 K 60 9170 K 140

950^810 1760^1560 3930^3720 6490^6270 9750^8990

2.14 K 0.15 1.82 K 0.29 2.62 2.60 K 0.18 328.17 K 0.10

6080 K 1130 11080 K 240 10960 K 275 11375 K 260 12035 K 810

8110^3680 11720^10470 11555^10425 12285^10595 13145^10635

n.d. n.d. n.d. n.d. n.d.

11405 K 150 11635 K 280

11925^11235 12115^11255

26.36 K 0.1 31.20 K 0.14

1805 K 70 7375 K 60 9235 K 205 9640 K 190 9725 K 135 10655 K 80

1885^1705 7435^7285 9760^8980 9930^9250 10280^9250 11045^10505

329.66 K 0.17 328.96 K 0.15 1.03 34.68 K 0.14 0.98 328.10 K 0.08

930 K 30 2225 K 75

1010^870 2320^2120

32.44 31.19

2660 K 50 7965 K 45

2740^2490 8050^7860

33.79 30.01

2450 K 50 7305 K 65

2590^2350 7420^7200

31.11 30.62

6345 K 55 6950 K 60

6450^6260 7090^6840

32.83 30.82

The radiocarbon ages were converted into calendar ages using the calibration programme Calib 4.3 (data sets INTCAL.98 and MARINE.98) after Stuiver et al., 1998. KI = samples conventionally dated at the Leibniz Laboratory (Kiel, Germany). KIA = samples AMS dated at the Leibniz Laboratory (Kiel, Germany). n.d. = not determined. a Midpoint of 1c range K half width of this range in calibrated years. True age of sample between these extremes for 2c probability. b Please, note that for the AMS dates N13 C includes the fractionation occurring in the sample preparation as well as in the AMS measurement and therefore are less signi¢cant for the original sample as compared to the conventional technique. c Samples dated with the conventional 14 C-method at the Laboratoire de Ge¤ologie du Quaternaire in Marseille (France).

N13 C of the gas was determined mass spectrometrically. The conventional radiocarbon ages are calculated with the 5568 years half-life (Stuiver and

Polach, 1977). The radiocarbon ages were all converted to calendar years before present (AD 1950) using the programme CALIB 4.3 (Stuiver et al., 1998). Marine samples were calibrated with the

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Table 2 Sea-level data (1c-range) as derived from the barrier reef core PMI-7, the in situ foreslope samples (fsl), the fringing reef cores C4, C5 and C8 and the sediment cores 89022, 89026, 89027 and 89028 Core number

Laboratory ID-number

Depth in core (m)

Palaeodepth estimated (m)

DSL (m)

Cal agea (yr BP)

PMI-7b C4c PMI-7b C8c PMI-7b C5c C4c PMI-7b PMI-7b PMI-7b PMI-7b PMI-7b 89022d 89028d 89028c 89028d 89026c 89026c 89026c 89027d 89027d 89026c in situ fslb in situ fslb in situ fslb in situ fslb in situ fslb

B.R.G.M. 0022 KI 4282, 021 B.R.G.M. 0093 KI 4385, 021 B.R.G.M. 0095 KI 4384, 021 KI 4383, 021 B.R.G.M. 0023 B.R.G.M. 0096 B.R.G.M. 0097 B.R.G.M. 0018 B.R.G.M. 0098 KIA 11560 KIA 8739 KI 4332, 001 KIA 5206 LGQ 600 LGQ 599 LGQ 601 KIA 11558 KIA 11559 LGQ602 B.R.G.M. 0200 B.R.G.M. 0197 B.R.G.M.079 B.R.G.M. 0196 B.R.G.M. 071

0.35 2.50 2.45 7.00 6.30 8.00 10.00 8.45 10.6 15.7 17.7 20.05 2.80 2.73 2.75 2.97 3.65 3.62 4.49 3.8 4.95 5.50 -

2 0.5 2 0.5 2 3 1 2 2 2 2 2 0 2 2 0 0.5 1 0.5 0.5 0.5 0 15 15 15 15 25

1.35 K 1 2.20 K 0.5 3.45 K 1 6.80 K 0.5 7.30 K 1 9.00 K 1 9.00 K 1 9.45 K 1 11.60 K 1 16.70 K 1 18.70 K 1 21.05 K 1 28.80 K 0 34.73 K 2 34.75 K 2 39.97 K 0 59.15 K 0.5 58.50 K 1 60.00 K 0.5 59.30 K 0.5 60.50 K 0.5 61.50 K 0 102 K 10 110.5 K 7.5 110.5 K 7.5 117.5 K 7.5 130.0 K 15

1500 K 100 2660 K 50 3700 K 200 6950 K 60 7200 K 400 7305 K 65 7965 K 45 8200 K 200 8200 K 300 8600 K 300 9300 K 300 9600 K 400 9170 K 140 9640 K 190 9725 K 135 10655 K 80 10960 K 275 11080 K 240 11375 K 260 11405 K 150 11635 K 280 12035 K 810 13600 K 400 16900 K 400 18200 K 500 17100 K 400 18400 K 500

DSL, depth below present sea-level. B.R.G.M., Bureau de Recherches Ge¤ologiques et Minie'res Orle¤ans (France). a Midpoint of 1c range K half width of this range in calibrated years. True age of sample between these extremes for 2c probability. b U/Th dates. c 14 C conventional method. d 14 C AMS method.

marine calibration curve, where a standard reservoir age correction of 400 yr is included (MARINE.98 ; Stuiver et al., 1998). The organic samples were calibrated with the dendrochronological calibration curve (INTCAL.98; Stuiver et al., 1998). Uncertainties in calibrated ages are reported as the midpoint of the 1c range K the half widths of this range in calibrated years. The true age of the sample lie between these extremes for the 2c probability. Natural variations in atmospheric and oceanic 14 C levels, especially during the deglacial period and early part of the Holocene, lead to sometimes

very broad and/or broken calibrated age ranges. The so-called 14 C-age plateaus with constant ages are observed at 12.7, 10 and 9.5 14 C kyr BP. This could be caused by changes in the 14 C production (the dipole moment of the earth, solar wind magnetic properties), carbon cycle changes and possible geochemical contamination (Bard et al., 1990, 1996; Eisenhauer et al., 1993).

4. Core location and samples Four cores from the fringing reef and ¢ve sedi-

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187

Fig. 3. Corelogs of the fringing reef cores derived from the northeastern Longoni Lagoon. Core C8 is derived from the reef top, cores C4 and C5 from the upper foreslope and core C1 from a lower slope to basin position landwards. Lithology, fossil content and calendar ages yr BP (midpoint of 1c range K half width of this range in calibrated years; true age of sample between these extremes for 2c probability) are indicated (after Zinke, 2000).

ment cores from the lagoon of Mayotte reaching the Pleistocene/Holocene boundary provide the opportunity to reexamine and complete the postglacial sea-level curve as derived from the barrier reef core PMI-7 and in situ coral samples from the foreslopes (Table 2; Camoin et al., 1997; Dullo et al., 1998). The fringing reef cores were drilled at the northeastern coast within Longoni Bay (Fig. 1). Core recovery was 70% for cores C8 and C5, 53% for C4, and 20% for C1. The cores are detritusdominated because they are derived from a protected area at the inner shoreline of Longoni Bay with moderate to low-energy conditions and low storm severity (Fig. 3). Under such hydrodynamic conditions the proportion of bioconstructors is

generally less than that of loose sediment and the characteristic sediments are gravelly to sandy (Montaggioni, 1988; Cabioch et al., 1995; Camoin et al., 1997; Braithwaite et al., 2000). The fringing reefs settled directly on basaltic bedrock or volcanic alterites. Thin horizons of massive to robust branching corals alternate with bioclastic sand to gravel. The bioclastic sand to rubble contain coral blocks (20^30 cm in diameter) to coarse coral debris (5^10 cm in diameter), molluscs, Halimeda, alcyonarian spiculae and various foraminifera. The accretion history between the fringing reef cores varies due to their position within the reef in relation to site-speci¢c wave energy and sediment erosion. Core C8 is derived from the inner reef £at resting on a topographic high,

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J. Zinke et al. / Marine Geology 194 (2003) 181^196 MAY.KL.89028 -34 m

MAY.Kl.89026 -56 m

MAY.Kl.89027 -55 m

0

0

MAY.KL.89022 -26 m

0

0

1

1

MAY.Kl.89011 -42 m 0

260 ± 350 1805 ± 70

depth (m)

1

900 ± 30 1655 ± 55 2

3850 ± 40

3

6360 ± 60 9170 ± 140 3

3

depth (m)

10655 ± 80

2

2

9235 ± 205 9640 ± 190 9725 ± 135

3

6080 ± 1130

depth (m)

2

1

7375 ± 60

depth (m)

depth (m)

1

2

11080 ± 240 10960 ± 275 4

4

4

11405 ± 150

5

11635 ± 280

9450 ± 470

11375 ± 260 5

3

12035 ± 810

6

LITHOLOGY carbonate mud carbonate sand carbonate sand to gravel mixed terrigenous-carbonate mud terrigenous mud mangrove mud

FOSSILS bivalves gastropods pteropods pelagic foraminifera benthic foraminifera echinoids corals bryozoans plant remains concretion

paleosol

Fig. 4. Corelogs for the studied sediment cores from the lagoon of Mayotte. The sediment cores are derived from the southwestern Boue¤ni Lagoon (cores 89026, 89027 and 89028), the southern Dapani Lagoon (core 89022) and the northeastern Longoni Lagoon (core 89011). In all cores, the pre-Holocene basement consists of palaeosoils. Present water depth, lithology, fossil content and calendar ages yr BP (midpoint of 1c range K half width of this range in calibrated years; true age of sample between these extremes for 2c probability) are indicated in the legend (after Zinke, 2000).

where wave energy and sediment reworking is highest. This is indicated by a coral block at 2.15 m dated at 6.3 kyr cal BP, the date of which is too old in comparison with sediments at a similar core depth in C4 and the known sea-level history for the Mayotte Lagoon. Cores C4 and C5 are from the reef crest to upper foreslope and show continuous accretion. Core C5 shows a coarsening upward above 4.5 m (2.45 kyr cal BP) with a higher abundance of coral blocks and debris, while their abundance decreases in

core C4. This is probably related to enhanced shedding of reefal detritus in relation to reduced accommodation space at the reef crest (C4). A volcanic event showing trachytic pumice is present in core C5 at a depth of 7^7.8 m. The age of this event, 7 kyr cal BP, agrees with the age of an ash layer present in the barrier reef core and the sediment cores from the lagoon and serves as an independent time marker (Zinke et al., 2000). Core C1 was taken from a small channel in a more landward position. The base of this core dates

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at 2.2 kyr cal BP suggesting a time lag of several thousand years in sediment accumulation at this position in comparison to cores C4, C5, and C8. This most probably can be related to unfavourable energy conditions for sediment production and deposition. The average Holocene sedimentation rate is the highest in the basinal environment (C1) with 3.16 mm/yr, followed by 1.17 mm/yr at the fringing reef slopes (C4 and C5) and 1 mm/yr in the reef £at sediments (C8). The sediment cores are derived from the southwestern Boue¤ni Lagoon (cores 89026, 89027 and 89028), the southern Dapani Lagoon (core 89022) and the northeastern Longoni Lagoon (core 89011) (Fig. 4). In all cores, the pre-Holocene basement consists of palaeosols (Zinke et al., 2001). Core recovery was 100% and depth in core was corrected for vertical compaction. Depth of sample relative to mean sea-level was corrected for tide level at the time of sampling. We examined sedimentary facies development and changes in faunal composition to infer palaeodepth for a speci¢c environment (Zinke, 2000; Zinke et al., 2001). The oldest £ooding level is preserved in cores 89026 and 89027 at a depth of 60^61 mbpsl (Elmoutaki et al., 1992; Zinke et al., 2001). Mangrove deposits, which contain in situ root fragments of Rhizophora and other plant remains, overlie an organic-rich palaeosol in both cores (Elmoutaki et al., 1992). Mangroves are excellent sea-level indicators, since they mark the uppermost level of the intertidal £at at about K 0.5 m. Core 89028, in which terrigenous muds overlie an organic-rich ferralitic palaeosol, provides a detailed look on the £ooding of the lagoon at 37 mbpsl. Two samples (one large oyster and one large coral block) from this level revealed consistent ages, 9.7 and 9.6 kyr cal BP, respectively (Fig. 4). Organic root fragments within the underlying palaeosol provide a signi¢cantly older age, 10.7 kyr cal BP. Core 89022 consists of coral^ mollusc sand to gravel overlying a ferralitic palaeosol at 29 mbpsl. Datings within the ferralitic soil (exposure horizon) and the ¢rst marine sediments show the presence of a time lag of several thousand years in sediment accumulation that most likely is caused by reworking and erosion. Time and depth of the £ooding level within core

189

89022 agrees with the dates found for the other cores. Most of the dates within the £ooding horizons are derived from organic samples, where no reservoir age correction must be applied. Therefore, the dates are of high quality. Core 89011 consists of coral^mollusc muddy sand overlying a ferralitic palaeosol. It provided additional data of the £ooding history of the northeastern Longoni Lagoon (Elmoutaki et al., 1992). The estimation of palaeodepth for a dated coral sample (Acropora sp.) deposited 20 cm above the £ooding horizon is problematic, while no further dating is available for this core. The sediments at this position in the core were interpreted as beach barrier or backreef deposits in proximity to the coastline (Elmoutaki et al., 1992). The radiocarbon ages in the sediment cores show an undisturbed depth^ age relationship suggesting that bioturbation only had minor in£uence on the overall sedimentation pattern (Table 1), which was probably caused by the high sedimentation rates. We chose signi¢cant changes in the slope of the sea-level curve as anchorpoints to calculate individual rates in sea-level variations (Fig. 5).

5. Results and discussion The early postglacial sea-level history for Mayotte for the time interval between 18.4 and 11.6 kyr cal BP bears some uncertainties (Dullo et al., 1998). The proposed sea-level position of 145 mbpsl at 18.4 kyr cal BP (Colonna et al., 1996) deviates signi¢cantly from other observations, i.e. 111^116 mbpsl (Hanebuth et al., 2000). The most simple explanation for the large di¡erence between this date and all other known sea-level stands at 18.4 kyr cal BP would be an increased living depth of the dated coral (Acropora sp.). The sea-level position of about 105 to 118 mbpsl between 16 and 18.4 kyr cal BP at Mayotte is also observed in other deglacial sealevel records (Fairbanks, 1989; Hanebuth et al., 2000; Yokoyama et al., 2000, 2001) (Figs. 5^7). A dated in situ massive coral Porites from the Mayotte foreslope of the barrier reef at 105 mbpsl gives an age of 13.6 K 0.4 kyr cal BP. The water depth range in which massive Porites sp. grows

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cal age yr BP 0

5

10

0

15

20

3 ± 1 mm/yr 8-7 kyr cal BP

0.9 mm/yr 7-1.5 kyr cal BP

9 ± 2 mm/yr 9.6-8 kyr cal BP

30

water depth (m)

19 ± 5 mm/yr 11.6-9.6 kyr cal BP (MWP-1B) Younger Dryas (12.7-11.6 kyr cal BP)

60

17.5 ± 10 mm/yr 13.6-11.6 kyr cal BP (MWP-1A)

90

120 5 mm/yr 18.2-13.6 kyr cal BP

150 0

5

10

15

20

cal age yr BP

cal age yr BP with error bars marine samples organic samples uncertain sea-level curve Fig. 5. Reconstructed sea-level curve on the basis of examination of in situ coral samples from the foreslopes of the barrier reef, a barrier reef core, four fringing reef cores and four sediment cores from the lagoon of Mayotte. Average rates of sea-level rise are indicated for speci¢c time intervals. Horizontal bars for each sampling point indicate the 2c-range of calibrated 14 C and/or U/Th dates, respectively. The vertical bars indicate the uncertainty in estimated palaeowater depths.

varies between 5 and 20 m (Montaggioni and Faure, 1997), but in general it occurs deeper than the aforementioned Acropora. Thus, taking into account a greater palaeo-water depth for this coral, this level would be consistent with sea-level estimates for this time interval from Barbados

(Fairbanks, 1989), the Caribbean (Chappell and Polach, 1991; Edwards et al., 1993), Tahiti (Bard et al., 1996), Hawaii (Fletcher and Sherman, 1995), the Sunda Shelf (Hanebuth et al., 2000) and Papua New Guinea (Blanchon and Shaw, 1995), which range between 64 and 85

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191

Fig. 6. Flooding history of the lagoon of Mayotte between 11.6 and 6 kyr cal BP showing the stepwise £ooding of the formerly subaerially exposed lagoon. The major passages are indicated by arrows.

mbpsl. U/Th data of all corals measured contain N234 U values equal to those of modern sea water and showed no evidence for any deviation from a closed system behaviour. Thus, the data given by our datings seem to be very reliable. Consequently, the assumption can be made that the depths and ages of the foreslope coral samples are valid. Therefore, we suggest that a rapid hydro^isostatic adjustment (submergence) of the small volcanic island of Mayotte with its surrounding oceanic crust due to large meltwater addition during the deglaciation might be a possible

mechanism causing such an age^depth discrepancy between the continental and larger island sea-level sites (Nakada and Lambeck, 1991; Lambeck, 1993; Grossman et al., 1998). While too many uncertainties remain in this respect in resolving this problem accurately, we will restrict our discussion to the period between 11.6 kyr cal BP to present, a period for which we have a large set of reliable data. The lagoon of Mayotte was exposed until 11.6 kyr cal BP. First £ooding of the lagoon occurred at about 11.6 kyr cal BP through the western

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0

0

5

10

15

20

15

20

-20

water depth (m)

-40

-60

-80

-100 Mayotte Tahiti Morley Sunda-Shelf Barbados

-120

-140 0

5

10 cal age yr BP

Fig. 7. Comparison of the sea-level curve of Mayotte with other Holocene sea-level curves. Note the o¡set between the sea-level curve for Mayotte and the others during the postglacial (19^11 kyr cal BP) and Middle to Late Holocene period (7^0 kyr cal BP). Good agreement is reached between 11 and 7 kyr cal BP.

passages in the Boue¤ni Lagoon. At that time, sea level reached a position of about 60 mbpsl (cores 89026 and 89027; Fig. 4). At ¢rst, only the deeper lagoonal areas in the northern and western lagoons deeper than 50 mbpsl were £ooded, while the northeastern and eastern lagoons remained exposed. This di¡erence is related to the depth of the passages through the Pleistocene barrier reef and the pre-Holocene bathymetry. The maximum depth of the passages at the western and southwestern lagoons lies at 70 mbpsl, while in the other lagoonal areas they do not exceed 40 mbpsl and even might not exceed 20 mbpsl (northeastern Longoni Lagoon, eastern Ajangoua

Lagoon, southeastern Saziley Lagoon and southern Dapani Lagoon) (Guilcher, 1965). The mangroves in cores 89026 and 89027 from the Boue¤ni Lagoon existed for about 500 years during a phase with declined rates of sea-level rise, which coincides with the Younger Dryas period (12.7^11.6 kyr cal BP). A decline to a rate of about 2 mm/yr was reported for the Papua New Guinea reef during the Younger Dryas (Edwards et al., 1993). Another indication for a decline in the rate of sea-level rise at this time is the presence of a coral reef terrace at 60 mbpsl found at the outer barrier reef slopes around the entire island (Dullo et al., 1998). The reefs were able to keep up with sea level during this speci¢c sea-level period, but drowned during the following rapid rise in sea level (Dullo et al., 1998). The mangroves also drowned after 11.6 kyr cal BP due to an abrupt rise in sea-level (cores 89026 and 89027; Fig. 4; Elmoutaki et al., 1992). The palynological association of the mangrove shows evidence for this sudden drowning event (Elmoutaki et al., 1992). The transition to the overlying carbonate muds is sharp, which is also shown by the geochemical data (Elmoutaki et al., 1992). The pollen spectra re£ect the development of non-brackish environmental conditions near the core site after 11 kyr cal BP. This drowning event is probably related to the so-called meltwater pulse 1B, which occurred around 11.3 kyr cal BP (Fairbanks, 1989, 1990; Bard et al., 1990, 1996; Chappell and Polach, 1991; Edwards et al., 1993). This sea-level rise for Mayotte Island after meltwater pulse 1B has an amplitude of about 23 mm/yr, comparable to rates reported in literature (Fairbanks, 1989; Bard et al., 1990, 1996; Edwards et al., 1993). After this speci¢c event sea-level rose continuously and reached a position of 37 mbpsl between 10.6 and 9.8 kyr cal BP (core 89028; Fig. 4) and of about 22 mbpsl at 9.6 kyr cal BP (Camoin et al., 1997; Dullo et al., 1998). Between 11.6 and 9.6 kyr cal BP the average rate of sea-level rise with approximately 19 K 5 mm/yr resulted in a total rise of about 38 m. First £ooding of the northeastern Longoni, eastern Ajangoua and southern Dapani lagoons occurred after 10 kyr cal BP. This is con¢rmed by a dating in core

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89011 from the Longoni Lagoon and core 89022 from the Dapani Lagoon (Fig. 4). Ages of 9.5 kyr cal BP can be given at 44 mbpsl in core 89011 and 9.7 kyr cal BP at 29 mbpsl in core 89022 (Fig. 4). Thus, £ooding of the eastern, northeastern and southern lagoons lagged behind the northern and western lagoons by approximately 1 kyr. The Holocene barrier reef started to grow at 9.6 K 0.4 kyr cal BP at a depth of 22 mbpsl within the eastern Ajangoua lagoon near Petite Terre Island (Camoin et al., 1997; Dullo et al., 1998). This date is consistent with the time of reef colonisation in western Australia (Eisenhauer et al., 1993) and other Indo^Paci¢c areas (Montaggioni, 1988), which occurred when sea level reached a position of 24 mbpsl before 9.8 kyr cal BP. The vertical reef growth rate for Mayotte was very high during the early stages of sea-level rise, reaching about 8.3 mm/yr (Camoin et al., 1997). Between 9.6 and 8 kyr cal BP, sea-level rose by about 9 K 2 mm/yr in average resulting in a total rise of 14 m reaching a position of about 8 mbpsl. This is con¢rmed by palaeobathymetric reconstructions of the barrier reef core PMI-7 and the fringing reef cores (Fig. 3; Camoin et al., 1997). The rate of sea-level rise between 9.8 and 8 kyr cal BP coincides with a small peak in the meltwater discharge curve from Barbados (Fairbanks, 1989) at about 9.5 kyr cal BP. A small peak in sea-level rise between 9 and 8 kyr cal BP is also present in the record from Abrolhos Island in western Australia (Eisenhauer et al., 1993). These peaks are not meltwater pulses, but most likely were caused by the redistribution of water masses for Barbados and delayed hydro^isostatic adjustment of the oceanic crust adjacent to the Australian shelf. Between 8 and 7 kyr cal BP, rates in sea-level rise declined to 3 mm/yr resulting in a total rise of 3 m and sea-level reached a position of approximately 5 mbpsl. After 7 kyr cal BP, sea-level rise slowed down to 0.9 mm/yr and reached its present position at about 2.5 kyr cal BP. Vertical reef accretion slowed down to 1.14 mm/yr after 7 kyr cal BP and stopped after stabilisation of sea-level, which is con¢rmed by the age of the reef top with 1.5 K 0.1 kyr cal BP (Camoin et al., 1997). The

193

maximum sedimentation rates in the sediment cores (0.8^1.2 mm/yr) and the growth rates of the barrier and fringing reef (0.91 mm/yr) lie in the range of the proposed rates in sea-level rise during this time interval. The datings of the fringing reef cores and lagoonal biogenic components in combination with several sedimentological events within the lagoonal cores allow us to specify the sequence of sealevel events within the southwestern Indian Ocean. The sequence of events within the lagoonal sediments and in the fringing reef cores agrees well with the scenario evolving from the barrier reefcore analysis for the interval between 9.6 kyr cal BP until present. This study completes the sealevel history for the time interval between 11.6 and 8 kyr cal BP. The preserved rates of sea-level rise in the lagoon of Mayotte represent average rates, including the meltwater pulse 1B, and coincide with the average rates observed in other deglacial and Holocene sea-level records (Fairbanks, 1989; Chappell and Polach, 1991; Blanchon and Shaw, 1995; Bard et al., 1996; Camoin et al., 1997; Dullo et al., 1998; Hanebuth et al., 2000). The greater depth range in the Mayotte sea-level curve and the transgressive shape may be related to the rapid isostatic adjustment of Mayotte and the overshoot of the sea-level rise relative to continental sites during sudden sea-level events. Small oceanic islands like Mayotte show a continuous rise because the addition of water changes the hydrostatic equilibrium and forces the island and the oceanic crust to move deeper into the mantle to compensate for the water load. Thus, the oceanic island record the full sea-level change, while the continental site records only a part of the total rise (Bloom, 1967; Eisenhauer et al., 1993). The di¡erence in the shape of the sealevel curves during the Middle to Late Holocene is probably also related to the ocean syphoning e¡ect (Mitrovica and Peltier, 1991; Grossman et al., 1998). Mayotte as a small volcanic island located in the ‘far-¢eld’ (far from former glaciated regions) shows a ‘transgressive’ sea-level curve (Camoin et al., 1997). Mayotte underwent no uplift after the termination of meltwater input into the ocean and the following redistribution of water masses towards the ‘near-¢eld’, because of

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its small size and intraplate volcanic origin between the East African continent mass and the microcontinent of Madagascar. Barbados as an oceanic island located in the forebulge region of the ‘near-¢eld’ (close to former glaciated regions) also shows a ‘transgressive’ sea-level curve, but with a deeper depth range. This is due to the redistribution of water masses towards the ‘near¢eld’ and the collapse of the forebulge (Mitrovica and Peltier, 1991; Eisenhauer et al., 1993; Grossman et al., 1998). The continental sites in the ‘far¢eld’ show a ‘transgressive^regressive’ sea-level curve because of isostatic compensation of the sea £oor adjacent to these regions and the redistribution of water masses to the ‘near-¢eld’ (Bloom, 1967; Camoin et al., 1997; Eisenhauer et al., 1993). Thus, the shape of the Middle to Late Holocene sea-level curves mirror the type of site choosen and their location in the ‘near¢eld’ or ‘far-¢eld’. We suggest that the aforementioned results from other studies and those from Mayotte support the idea that the ocean syphoning process is the major cause of Late Holocene ‘transgressive’ sea-level curves close to former glaciated regions.

6. Conclusions The postglacial sea-level history between 18.2 kyr cal BP until present was reexamined and completed on the basis of cores from the fringing reef and the lagoon of Mayotte. The rates of sea-level rise in the lagoon of Mayotte represent average rates, and also include meltwater pulses 1A and 1B. It revealed a sequence of events characterised by a sea-level rise at a rate of: (1) about 19 mm/yr between 11.6 and 9.6 kyr cal BP, (2) 9 mm/yr between 9.6 and 8 kyr cal BP, (3) 3 mm/yr between 8 and 7 kyr cal BP, and (4) 0.9 mm/yr after 7 kyr cal BP until stabilisation at the present level at 2.5 kyr cal BP. A decline in the rate of sea-level rise or even a stillstand is observed before 11.6 kyr cal BP, which coincides with the Younger Dryas. The sea-level rates between 11 and 8 kyr cal BP show an almost linear transgression with a prominent change in the inclination of the sealevel curve between 8 and 7 kyr cal BP.

The £ooding of Mayotte Island was controlled by the depth of the reefal passages, which originate from erosion by rivers and/or karstic processes during the time of emergence since the last interglacial (isotope stage 5a, 80^60 kyr cal BP). Flooding of the lagoon occurred at 11.6 kyr cal BP through the deepest passages (60^70 mbpsl) in the northern, western and southwestern lagoons. It preceded the £ooding of the other lagoons by approximately 1 kyr. The sequence of events within the lagoonal sediments and in the fringing reef cores agrees well with the scenario evolving from the barrier reef-core analysis for the interval between 9.6 kyr cal BP until present. This study completes the sealevel history for the time interval between 11.6 kyr cal BP and 8 kyr cal BP. The observed average rates in sea-level rise agree well with the average rates observed in other deglacial and Holocene sea-level records (Fairbanks, 1989; Chappell and Polach, 1991; Blanchon and Shaw, 1995; Bard et al., 1996; Camoin et al., 1997; Dullo et al., 1998; Hanebuth et al., 2000; Yokoyama et al., 2000, 2001). Di¡erences in the shape of the sea-level curve from Mayotte to other sites located far from former glaciated regions may be related to: (1) the small size of the island, (2) the rapid downward movement of this small volcanic island with the oceanic plate into the mantle due to hydro^ isostatic compensation after addition of water masses, and (3) the location between large islands and continent masses.

Acknowledgements The gravity cores from the Mayotte Lagoon were provided by the French research programme CORDET 1989^1990 during the cruise CARLAMAY with the scienti¢c vessel N.O. La Curieuse in 1989 by researchers from the University of AixMarseille and the University of La Re¤union. The fringing reef cores were drilled by the Geotechnical Service in Mayotte for the ‘Chambre Interprofessionelle’ of Mayotte in 1997. We thank Mr. P. Andrieux for providing us the opportunity to sample and study the cores. Financial support for this study was provided by the German Sci-

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ence Foundation (DFG-Grant DU129/14 to W.D., J.R, and J.Z.). We would like to thank Dr. Frank Bruhn and the technicians of the Leibniz-Labor at the Christian-Albrechts-University in Kiel (Germany) for the radiocarbon datings. We also thank the ‘Laboratoire de Ge¤ologie du Quaternaire Marseille’ (France) for providing radiocarbon datings and the laboratory B.R.G.M.Orle¤ans (France) for the U/Th analysis. This research is part of the geomorphological programme of the G.I.S. ‘Lag^May’ (Marine and littoral environments of the island of Mayotte) supported by the ‘Conseil Ge¤ne¤ral de Mayotte’. We would also like to thank Anton Eisenhauer and Till Hanebuth (both at GEOMAR, Kiel) for their constructive reviews of an earlier version of the manuscript.

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