BiologicalJournal of the Linnean Society (109~5),56: 13-42. With 22 figures
m e Pitcairn Islands: biogeography, ecology and prehistory
Edited by T. G. Benton and T. Spencer
The Pitcairn Islands, South Pacific Ocean:
plate tectonic and climatic contexts
T. SPENCER
Department of Geography, University of Cambridge, Downing Place,
Cambridge CB2 3EN
The remote Pitcairn Group in the South Pacific Ocean comprises a volcanic island (Pitcairn
Island), two low coral atolls (Oeno, Ducie) and a raised coralline island (Henderson Island).
The geological history of these islands, on anomalously thin oceanic lithosphere, is related
to the development of two subparallel island chains (Oeno-Henderson-Ducie; Pitcairn)
associated with intra-Pacific plate ‘hotspot’ activity; the surface manifestation of this activity
has been partly determined by structural lineations in the plate inherited from past plate
history. The climate of the Pitcairn Islands is determined by the position of the subtropical
high pressure system and the South Pacific Convergence Zone. Variations in the strength of
this atmospheric circulation system, measured by changes in the Southern Oscillation index
of pressure difference, provide a partial explanation of the long-term variability of mean
annual rainfall at Pitcairn Island. Knowledge of past climates in the Pitcairn Group remains
speculative. Maps of the Pitcairn Islands and a report of climate at Henderson Island (2/911/92) are included in the paper.
0 199.5 The Linnean Society of Landon
ADDITIONAL KEY WORDS:-hot spot - atoll
Pacific Convergence Zone - palaeoclimate.
-
seamount
-
Southern Oscillation
-
South
CONTENTS
Introduction . . . . . . . . . . . . . . . .
Location . . . . . . . . . . . . . . . .
Plate tectonic contexts . . . . . . . . . . . . . .
Theoretical and technical advances in understanding Pacific island histories
The South Pacific ‘superswell’ . . . . . . . . . . .
Plate geodynamics and island histories in the Pitcairn Group . . .
Climatic contexts . . . . . . . . . . . . . . .
Characteristics of the South Pacific Ocean climate system . . . .
The Southern Oscillation and the South Pacific Convergence Zone . .
ENSO episodes and rainfall at Pitcairn Island . . . . . . .
Palaeoclimates of the Pitcairn Islands: Some speculations . . . .
Acknowledgements
. . . . . . . . . . . . . .
. . . . . . . . . . . . . . . .
References
Appendix A: Maps of the Pitcairn Islands . . . . . . . . .
Appendix B: The climate of Henderson Island 1991-1992 . . . . .
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T. SPENCER
INTRODUCTION
Location
The islands of the Pitcairn Group-volcanic Pitcairn Island, the small coral
islands of Ducie and Oeno Atolls and the raised coralline Henderson Islandlie between 23.9"S-24.7"S and 124.7"W-130.7"W in the 75 Mkm2 of the
South Pacific Ocean. The islands are isolated from their nearest
neighbouring island groups-the most easterly island, Ducie Atoll, is 1000 km
west of Easter Island whilst the most northerly and westerly, Oeno Atoll, is
450 km east of the Minerve reefs and the Gambier Islands-and
are
extremely remote in a Pacific Ocean basin context: New Zealand and South
America are both over 4500 km distant (Fig. 1). This paper, as an introduction
to the symposium volume, reviews the geological history of the Pitcairn
Islands in a long-term and large-scale plate tectonic context and analyses
historical and longer timescale climatic variability in this area within the
framework of the complex dynamics of the Pacific atmosphere-ocean system.
Maps of the four islands, some the result of new surveys, are presented in
Appendix A. Information on climatic conditions at Henderson and Pitcairn
while the Sir Peter Scott Commemorative Expedition was in the field, and
their long-term representativeness, is provided in Appendix B.
-
11O"W
Figure 1. Plate boundary (East Pacific Rise), island chains and structural lineations (FZ = fracture
zone), and location of the Pitcairn Group, South Pacific Ocean. Bathymetric contours in
metres.
GEOLOGICAL AND CLIMATIC BACKGROUND. PITCAIRN GROUP
15
PLATE TECTONIC CONTEXTS
Theoretical and technical advances in understanding PaczJic island histories
Pacific island histories can only be satisfactorily explained by reference to
the characteristics of the plate which underlies them. Until recently, however,
the submarine structure and topography of the Pacific plate was poorly
known, with ocean floor bathymetry mapped by irregular networks of 6-26
(at most 50) shiptracks per Mkm2. In recent decades, however, ocean floor
mapping has been revolutionized by the development of satellite altimetry
techniques; these techniques permit mapping of large areas at typical densities
in the Pacific Ocean of 36 transits per Mkm' near the equator, improving
to 70 transits per Mkm' at higher latitudes where track spacing decreases to
less than 40 km (Mayes, Lawyer & Sandwell, 1990). Such high resolution
mapping allows both detection of previously unknown structures and
speculation as to their plate tectonic significance. Subsequent multi- and widebeam sonar mapping and submersible observations can then be used to
determine the fine detail of these structures. It is the combination of these
technical advances with the chronology derived from seafloor magnetic
lineations that provides the basis for improved interpretations of the geological
history of the South Pacific Ocean, including the region of the Pitcairn
Islands.
The present Pacific plate occupies two thirds of the ocean floor south of
and Nazca plates to the east by
the equator. It is separated from the COCOS
the spreading ridge of the East Pacific Rise and from the Antarctic plate by
the convergent Pacific-Antarctic Rise. Oceanic lithosphere is consumed in
the South Pacific region at the Tonga-Kermadec trench system. The seafloor
spreading that has produced the present SW Pacific began at -84 Myr BP,
with a NNW spreading direction being replaced by a WNW trend at -42
Myr BP (Mayes et al., 1990). Present plate migration rates, established from
K-Ar dating of exposed island and atoll basement volcanics, are discussed
in more detail below; they average 10-11 cm a p l (Duncan & Clague, 1985).
These lateral movements have been, and continue to be, accompanied by
important changes in plate characteristics: the oceanic lithosphere cools,
thickens and subsides as it moves away from the spreading ridge. Parsons
& Sclater (1977) showed that an exponential relationship exists between plate
age and water depth, as modelled by a simple heat-loss equation. Thus any
physiographic feature formed at a mid-ocean spreading ridge (including
volcanoes) will subside with the plate. In the case of features above sea
level, coral growth is likely to accompany this subsidence where ocean waters
exceed a mean monthly temperature of -20°C in the coldest month of the
year. Continued upward coral growth as island subsidence proceeds may
lead to a volcanic island with a fringing reef being transformed into a barrier
reef island and, ultimately, a coral atoll (Darwin, 1842). This has been the
history of the coral atolls of the Tuamotu plateau, NW of the Pitcairn Group,
which have developed over volcanic basements originally formed at 45-30
Myr BP on or near the Pacific Ocean spreading centre (Mayes et al., 1990).
Studies of plate loading by volcanoes confirm that the elastic thickness of
16
T. SPENCER
the lithosphere, Te, increases with the age of the plate at time of loading:
thus 59 Myr old lithosphere has a typical Te of 19 km whereas 80 Myr old
lithosphere is characterized by Te values of 25-30 km (Watts & Ribe, 1984).
These differences can be used to indicate whether islands were formed at
mid-ocean spreading ridges on this plate or at mid-plate locations characterized
by much thicker lithosphere. Variation in the height of the sea surface-the
marine geoid-reveals mass deficits and surpluses associated with seafloor
topography when appropriately filtered (e.g. Sandwell, 1984). Subtraction of
actual or modelled seafloor bathymetry from this geoidal signal results in a
map of geoidal anomalies. This information can then be used to indicate
island origins: spreading-ridge genesis is indicated by small amplitude (0.40.5 m per km of seamount height), short wavelength anomalies whereas offridge origin is indicated by large amplitude (1.4-1.5 m per km), large
wavelength anomalies (Watts & Ribe, 1984). Application of these principles
to the more southerly Tuamotu atolls (Okal & Cazenave, 1985) and the
island chains of the South Pacific-the Austral-southern Cook Islands, the
Society Islands and the Pitcairn Island-Gambier Islands lineation (Fig. 1)suggests an off-ridge origin at intra-plate thermal anomalies, or 'hotspots'.
Seafloor thermal swells, seismic swarms, hydrothermal activity and active
magma accumulations currently characterize the south-eastern extremities of
these chains (e.g. Macdonald seamount, Austral Islands: Stoffers et a l , 1989;
Mehetia, Society Islands: Binard et a l , 1993; (the Pitcairn hotspot is discussed
in more detail below)). In time, with continued plate motion, the lithosphere
becomes de-coupled from the fixed heat source below and the plate coolingseafloor sinking relationship is re-imposed, although at accelerated rates of
subsidence as plate re-heating imparts the characteristics of young lithosphere
(Detrick & Crough, 1978). As islands form, subside and disappear, so the
temporal Darwinian sequence of coral island types (i.e. fringing reef-barrier
reef-atoll) should be seen over space, also aligned in the direction of plate
motion (Stoddart, 1973). However, the admixture of island types in less than
perfect sequence within most of these South Pacific island groups and the
presence of raised reef limestone, or Makatea, islands, in close juxtaposition
to other island types, suggest that many island chains have more complex
histories (Fig. 2; Menard, 1986).
The Pitcairn Group, with its volcanic island, two atolls and raised coralline
island, represents a microcosm of these complexities. In addition, improved
bathymetry for the Group acquired from sea surface and sea bed remote
sensing technologies shows that further information must be added to the
simple pattern of the four emergent islands: Oeno Atoll rises from the
southern side of a broad plateau at -1600 m and Ducie Atoll is not a
single feature (Mammerickx et a l , 1975) but the surface expression of a field
of seamounts (Canadian Hydrographic Service, 1982). Further to the east,
clearly visible in the pattern of geoidal anomalies, is a major structure topped
by two seamounts reaching 1000 m below present sea level at 25.0°S,
122.2"W and 24.8"S, 121.7"W; Okal (1984) has proposed that this feature be
known as Crough Seamount. Geoidal anomalies also reveal additional
seamounts around 25.6"S, 121.2"W and 26.2"S, 121.8"W and near the East
Pacific Rise (Okal & Cazenave, 1985). How can these structures, and their
arrangement, be explained?
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
I
I
I
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A
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Fatu Hiva
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Rurutu
Mangaia
A A
-"
o
Tubuai
Rimatara
D
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Mangareva
0
0
160"W
0-1.5 my volcano
Older volcano
Uplifted island
Atoll
m
L"
MururoaiO
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0 .
Huka
Hiva Oa
8
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MA ROUESAS
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A
A
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0
0
0
Ai t uta ki
I
*tAUa
0
0
17
Henderson
%
O
Pitcairn A
%s
A
Duel;
Rapa
A
Macdonald
seamount
150"W
140"W
130"W
I
I
I
30"s -
f i e South Pactjic 'superswell'
One particular characteristic of the area of the South Pacific which includes
the Pitcairn Group is the fact that ocean floor subsidence accompanying
seafloor spreading and plate migration appears to have taken place more
slowly than expected (Cochran, 1986). An area of 18 Mkm2 exhibits water
depths 250-750 m shallower than predicted by Parson & Sclater's (1977)
plate cooling model. This region, 2000 km N-S by 5000 km E-W, and
bounded to the east by the spreading ridge of the East Pacific Rise, to the
south by the island chain of the Cook-Austral Islands and to the north by
the Marquesas fracture zone, has been termed the South Pacific 'superswell'
by McNutt and her co-workers (McNutt & Fischer, 1987; McNutt &Judge,
1990). The shallow bathymetry can be modelled if the temperature structure
of the plate is based on a thermal thickness of 75 km and a basal temperature
of 1385"C, rather than the standard figures of 125 km and -1350°C (Fig.
3; McNutt & Judge, 1990). These authors have argued that it is enhanced
heat flow and a low viscosity zone beneath the lithosphere in this area which
prevent the development of expected lithosphere thicknesses and limit the
plate to a maximum thermal thickness of 75 km; the shallow depth might
be maintained by small-scale thermal instabilities, organised into convective
rolls by the shear imparted from a fast-moving plate, at the base of the
plate. These convective processes may explain the 2000 km long gravity
lineations which run approximately parallel to the direction of plate motion
(Haxby & Weissel, 1983; Goodwillie & Parsons, 1992). McNutt & Judge
(1990) went on to argue that the thinned and less viscous lithosphere already
existed when hotspot volcanism began and thereafter was easily penetrated
T. SPENCER
18
-E
3000
2500
Tuarnotu Islands
-
- 125-krn-thick plate
-
- 75-krn-thick plate -.- Tuarnotus removed
-
I
1
-5001
I
I"
I
Figure 3. Comparison of depth profile over the South Pacific superswell referenced to plates
125 km and 7Fi km thick and with/without the Tuamotu Plateau removed. After McNutt &
Judge (1990).
by numerous, but often weak, thermal plumes. Thirty percent of the heat
flux of all the Earth's hotspots is liberated in this region (which covers only
3010 of the globe) and recent analyses of SeaBeam bathymetry transits have
shown that there is a threefold increase in seamount abundance south of the
Marquesas fracture zone, with abundances being equally high on young (018 Myr) and old (20-60 Myr) lithosphere (Bemis & Smith, 1993).
These arguments have been supported by the use of satellite altimetry
profiles to predict the lithospheric thicknesses supporting hotspot volcanoes
located on the superswell (Calmant, 1987; Calmant & Cazenave 1986, 1987).
If the lithosphere surrounding hotspot volcanoes has the thermal structure of
125 km thick plate then the Te values should correspond to the 40O+2O0C
isotherm. In fact, the estimates of Te for the superswell volcanoes appear 510 km less than expected, corresponding to the depth to the 400°C isotherm
of a plate less than 50 km in thickness (Fig. 4A).
However, in recent years this superswell model has been challenged
through two lines of argument. Firstly, the average heat flow for the superswell
does not differ dramatically from that of lithosphere of similar age elsewhere
on the Pacific plate and neither requires nor excludes the possibility of
thinner than 125 km plate. The heat flow datasets are, however, at variance
with the idea of a plate with a thermal thickness of less than 60 km (Stein
& Abbott, 1991). Secondly, calculations of Te values, which involve forward
modelling and calibration of model outputs, are heavily dependent upon
good quality bathymetric datasets. The recent availability of much improved
bathymetric databases has allowed re-calculation of Te values for different
island group hotspots (Fischer, McNutt & Shure, 1986; Goodwillie & Watts,
1993; Filmer, McNutt & Wolfe, 1993); these are detailed in Fig. 4B. At
Maupiti Island, Macdonald seamount, Maria Atoll and Pitcairn Island, these
revisions confirm previous analyses or estimate even lower elastic thicknesses.
However, more normal Te values are estimated for the Gambier Islands and
southern Cook Islands (Aitutaki, Rarotonga) and confirmed for Tahiti and
the Marquesas. This suggests that the area of anomalous plate thickness
values is more localized than previously thought (Goodwillie & Watts, 1993).
Furthermore, an alternative explanation for the characteristics of this region
is suggested: that intrusive volcanism has exploited an area of the Pacific
plate mechanically weakened by the fracturing, flexing and reheating of a
lithosphere affected by a complex tectonic history.
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
19
a) Previous T, estimates
-E
O -
Y
v
t-"
-5-10 -15 -
-20 -25 -30
-30
..
I
0
10
20
I
1
I
30
40
50
60
70
80
90
100
30
40
50
60
70
80
90
100
b) Revised T, estimates
-15
--
.
0
I0
20
Age of oceanic lithosphere at time of loading (Myrs)
Figure 4. Elastic thickness values for south central Pacific seamounts. Superimposed isotherms,
in "C, from standard plate cooling model; 450°C isotherm characteristically marks the base
of the elastic lithosphere. (a) Previous estimates of elastic thickness. (b) Revised values of
elastic thickness using high-resolution altimetry. Locations: P = Pitcairn; G = Gambier Is.;
MD = MacDonald seamount; MA = Maria, Austral Is.; MQ= Marquesas; MP = Maupiti;
A = Aitutaki;
RO = Rarotonga; T = southeastern Society Is.; RT = Raiatea-Tahaa;
H = Huahine. After Goodwillie & Watts (1993).
Plate geodynamics and island histories in the Pitcairn Group
The history of the Pacific plate in the SE Pacific Ocean since -26 Myr
BP, when the Pacific-Farallon ridge collided with the North American plate,
has been a series of clockwise-spreading ridge re-orientations and 'jumps' of
active ridge position (for full reviews: Handschumacher, 1976; Williams,
1986; Spencer, 1989; Mayes et al., 1990). These processes are still current:
at both 22-27"s and 31-35"s the East Pacific Rise is in the process of
'jumping' from a western to an eastern ridge, and thus defining the Easter
and Juan Fernandez micro-plates respectively (Schilling et al., 1985; AndersonFontana et al., 1986). A great variety of distinct structural lineations have
20
T. SPENCER
Pacific Plate
Figure 5 . Association of Oeno, Henderson, Ducie Islands and Crough seamount with Fracture
Zone 2 and independent lineation of Pitcairn-Gambier Islands. After Okal & Cazenave
(1985).
been associated with these and other geodynamic processes: (1) N170"E, a
former spreading ridge orientation; (2) N 150"E, an apparently old hotspot
trace south of the Austral Islands; (3) N80"E, the orientation of the Austral
and Marquesas fracture zones, former spreading ridge transform faults; (4)
N95"E, fracture zones FZ 1 and FZ2, re-orientated and re-located transform
faults related to (2) and more extensive lineations west of Pitcairn Island; (5)
NllO'E, orientation of the transform faults of the current East Pacific Rise
spreading centre and; (6) N120"E, the lineation of the contemporary island
chains of the SE Pacific (Diament & Baudry, 1987). These structures are
seen not only at the large-scale but also in the orientation of radial eruptive
fissures, or 'rift zones', on individual seamounts (e.g. Binard et al, 1992). In
the Pitcairn region, these structural lineations are associated with hotspot
activity and are seen in two groupings: one involving Pitcairn Island and
one concerned with a Crough-Ducie-Henderson-Oeno alignment (Fig. fi; Okal
& Cazenave, 1985; Spencer, 1989).
Pitcairn Island, 4 x 2 km and reaching 347 m above present sea level, is
the youngest subaerial expression of a 1900 km-long hotspot lineation,
orientated N120"E, which extends through the Gambier Islands and Mururoa
atoll, south-eastern Tuamotu archipelago to the Duke of Gloucester Islands.
The active hotspot region has recently been located on -30 Myr old oceanic
crust in a zone 40 to 110 km southeast of Pitcairn Island. At 3750 m, the
seafloor in the hotspot region is 750 m shallower than expected for plate of
this age and is characterized by the N170"E and N80"E structural lineations.
Volcanic activity is indicated by a field of over 20 volcanic hills; these have
been studied in detail by side-scan sonar, underwater television, coring,
dredging and water sampling techniques. As in the Society and Austral
hotspot regions (Binard et a l , 1992), the volcanoes are a mixture of (1) many
small (<500 m tall), often flat-topped hills, which are probably a mixture of
inherited features from former spreading ridge volcanic activity (as suggested
by Fonari, Batiza & Luckman, 1987) and more recent forms, and (2) a few
large (>500 m high) cones. In the latter category, two large conical
seamounts, named 'Adams' and 'Bounty' by Binard, Hekinian & Stoffers
(1992), reach to within 55 and 450 m of sea level respectively and show
evidence of very recent lava flows and hydrothermal activity (Stoffers et al,
1990).
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
21
The geology of Pitcairn Island has been resolved into four volcanic
formations related to a single shield volcano (Carter, 1967; Woodhead &
McCulloch, 1989). Overlying the lava flows of the Tedside Volcanics, the
general morphology of the island is dominated by the outcrop of the
Christian Cave Formation, a semi-circular caldera structure of trachytes,
andesites and tuffaceous materials, open to the north east. Associated with
this tuff-ring are the basaltic Pulawane Volcanics, which form the westernmost
part of the island, and the Adamstown Volcanics which fill and spill over
the caldera to the north and east. Potassium-argon dating of these volcanics
has identified two main phases of volcanism, at 0.95-0.76 and 0.63-0.45
Myr BP, with the older period probably representing the main phase of
island construction (Duncan et a l , 1974). 160 km to the west of Pitcairn
Island, the basaltic lavas which form the islands of Mangareva, Aukena and
Makapu, Gambier Islands, cooled between 7.1 and ,5.3 Myr BP (Brousse et
a l , 1972) and recovered basalts from Fangataufa Atoll and Mururoa Atoll
have been dated to 12.9-9.6 Myr BP and 11.8-10.7 Myr BP respectively
(Guillou et a l , 1993; Gillot, Cornette & Guille, 1992). Plate migration rates
on this lineation have been calculated at 12.7f 15.5 cm a-l (Duncan &
Clague, 1985) for the Pitcairn-Gambier sequence and at 10.7-11.0 cm a c l
(Brousse, 1985) when extended to Mururoa Atoll. These propagation rates
are comparable to the values calculated for other South Pacific island chains
(e.g. Austral-Cook Islands: 10.7f 1.6 cm ax1; Marquesas: 10.4f 1.8; Society
Islands: 10.9k 1.0; Duncan & Clague, 1985).
The history of the Crough-Ducie-Henderson-Oeno lineation is more ancient
and more complex. At first glance, these islands appear to be an extension
of a hotspot trace from the Tuamotu atoll of Hao south-east to the Acteon
Group and Marutea; they exhibit strong geoidal anomalies, indicative of an
off-ridge origin for their basements. The extension of the Hao-Marutea trace
passes through the location of Oeno and predicts the present position of the
hotspot to be -300 km south of the south-western boundary of the Easter
micro-plate. Henderson Island, Ducie and Crough Seamount, do not, however,
fall on this alignment but form a linear chain at a 15" angle to it (Fig. 5).
It has been suggested (Okal & Cazenave, 1985), following the theory of
Morgan (1978) and the idea of a south Pacific 'hotline' linking melting
anomalies from the Tonga trench to the Nazca plate (Bonatti & Harrison,
1976; Bonatti et a l , 1977; Turner & Jarrard, 1982), that this deviation resulted
from the interaction of the hotspot trace with a line of lithospheric weakness
represented by an old fracture zone known as FZ2 and showing the N95"E
orientation. Lateral leakage of magma led to the successive construction of
the presently three subaerial and one submarine Pitcairn Group structures.
If the process of island formation remains speculative then the timing of
these processes remains even more poorly defined. Okal & Cazenave (1985)
give ocean floor age estimates of 10 Myr for Crough Seamount, 14 Myr for
Ducie, 19 Myr for Henderson and 27 Myr for Oeno, and have calculated
the age of the plate at the time of loading, or flexural age, as 5-7 Myr.
This would suggest the following island ages ( f 1 Myr):-Crough Seamount:
4 Myr; Ducie: 8 Myr; Henderson: 13 Myr; and Oeno: 16 Myr. These dates,
however, can only be seen as first approximations as the flexural age is not
a true measure of the age of the plate at time of loading because of the
T. SPENCER
22
likelihood of plate re-heating at the time of island emplacement (for theory:
Detrick & Crough, 1978; McNutt, 1984).
Cessation of volcanic activity and island subsidence subsequently led to
the development of carbonate caps, of unknown thickness, on the islands of
Oeno, Henderson and Ducie (and Crough?). With the later, and independent,
development of the Duke of Gloucester Islands-Pitcairn Island hotspot
lineation, Henderson was affected by lithospheric flexure processes. O n the
relatively thin and deformable lithosphere of the South Pacific, the
emplacement of relatively young ( < 2 Myr) volcanoes, as a result of ‘hotspot’
activity, has produced a near-volcano moat and a peripheral bulge, or arch,
at some distance from each load. The coincidence of sea-level coral reefs
with this arch has resulted in the formation of raised reef islands (McNutt
& Menard, 1978). Flexural moats have been observed around some midplate volcanoes (e.g. Hawaii: Watts et aL, 1985; Marquesas: Fischer et a l , 1986)
and reef limestone uplift, apparently associated with flexure, demonstrated from
the Society Islands (Pirazzoli, 1983), the N.W. Tuamotu Archipelago (Lambeck,
1981a; Pirazzoli & Montaggioni, 1985) and the southern Cook Islands
(Lambeck, 1981b; Stoddart, Spencer & Scoffin, 1985; Spencer, Stoddart &
Woodroffe, 1987; Calmant & Cazenave, 1986). In the Pitcairn group
lithospheric flexure under the weight of Pitcairn Island has resulted in c.
30 m of uplift at Henderson. The raised reef topography of Henderson,
and estimates of the rate of flexure-controlled uplift, are considered in more
detail by Blake (1995).
CLIMATIC CONTEXTS
Characteristics of the South Pacijic Ocean climate system
The climate system of the South Pacific Ocean in general terms is
dominated by four linked features (Fig. 6; Streten & Zillman, 1984). First, a
180’
120‘w
30”
0’
30‘5
Figure 6. Mean streamline convergence for February over the Pacific Ocean, forming an
Intertropical Convergence (ITC) and South Pacific Convergence Zone (SPCZ). After Barry &
Chorley (1992).
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
23
belt of high pressure centred on latitudes 25-30"S, with a large, semipermanent anticyclone near longitudes 90-100"W. Second, south easterly
Trade Winds of generally moderate strength from the northern m a r p of
the high pressure area. An important feature of the Trade Winds is the
presence of high level return flows, with westerly winds above a strong
temperature inversion at 1200-2500 m. Third, the southern hemisphere
Trades meet those of the northern hemisphere at the Intertropical Convergence
Zone (ITCZ), a band of cumulonimbus cloud clusters and cloud-free areas
ranging from 3-10"N in the eastern central Pacific. In the western Pacific
Ocean, this doldrum belt lies within 5" of the equator, a region of fluctuating
light easterly winds, and, when associated with monsoonal systems in North
Australia in the austral summer, squally westerlies. These convergences are
often associated with high convective rainfall totals, particularly with the
addition of orographic influences over Indonesia and Papua New Guinea
and the SW Pacific (e.g. typically 5000 mm yr in Vanuatu: Taylor, 1973).
Although the ITCZ does not enter the southern hemisphere, it is associated
with the fourth unit of interest, the South Pacific Convergence Zone (SPCZ)
which extends from Papua New Guinea in a south easterly direction to c.
30"s 120"W and forms an important climatological link across the Tropics
and Subtropics. At its northern, low-latitude end, the SPCZ marks the
convergence of north easterlies associated with the ITCZ with south easterlies
ahead of high pressure systems moving eastwards from Australia and New
Zealand; its south easterly extension is caused by interactions with midlatitude westerlies; and its southern end is associated with wave disturbances
on the South Pacific Polar Front.
7 h e Southern Oscillation and the South PaczJic Convergence Zone
These components are part of a much greater ocean-atmosphere system
the influence of which extends well beyond the Pacific Ocean basin (Bjerknes,
1969). It is known as the 'Southern Oscillation' (the large-scale exchange of
atmospheric mass between the tropical and subtropical parts of the West
Pacific, South Pacific and Indian Oceans) or the 'Walker Circulation' (the
associated east-west vertical circulation cell). Variations in the strength of
this circulation system are measured by the Southern Oscillation Index (SOI),
a normalized measure of the difference in atmospheric pressure between
Tahiti, representing the subtropical high pressure system, and Darwin, North
Australia, measuring the strength of the equatorial low pressure system (Fig.
7). Furthermore, the index can also be used to identify ocean circulation
system characteristics which are linked to, and modulate, atmospheric
circulation changes. In extreme cases, high, positive values of the index, or
'cold phase' periods, are associated not only with strong trade winds and
high rainfall over northern Australia and SE Asia but also with strong, trade
wind-driven upwelling off the coast of South America and the presence of
anomalously cold surface waters along the equator. This leads to aridity on
low islands along the equator. When the SO1 index falls and the trade winds
slacken, warm water 'sloshes back' to the eastern Pacific (Wyrtki, 1971i). This
leads to a shift in convective activity to the east, with high rainfall on the
usually arid equatorial islands and the South American coast and drought in
24
T. SPENCER
Figure 7. Correlations of yearly pressure at Darwin with South Pacific stations. Note negative
correlation with Tahiti. Solid circles = 30 years of record (1944-73), open circles = incomplete
records. After Trenberth (1976).
the typically ever-wet northern Australia and SE Asia region (Fig. 8). At its
fullest development the warm phase leads to ‘El Nino’ conditions when
warm, nutrient-poor southward flowing surface currents replace the normally
cold, highly productive waters of the northward-moving, upwelling Humboldt
current along the coasts of Ecuador and Peru. Coldiwarm phase oscillations
Figure 8. Patterns of precipitation and temperature anomalies within the Pacific Ocean basin
and beyond associated with warm phases of the ENS0 cycle.
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
25
are quasi-periodic, occurring at 2- 10 year intervals, with a typical duration
of a full-scale oscillation being 12-18 months. Individual events, however,
vary considerably in their time of commencement and duration; strength of
development; and pattern of evolution (Philander, 1990). The turn-on and
turn-off of these events appear to be controlled by the re-distribution of
near-equatorial cold and warm water masses associated with the interactions
of long, westward-propagating Rossby waves and fast-moving Kelvin waves
travelling east (Cane, 1992); after particularly 'strong' warm phases such
redistributions may continue to affect the global climate system for as long
as a decade later (Jacobs et aZ., 1994).
Variations in the Southern Oscillation Index also lead to variations in the
location of the South Pacific Convergence Zone (SPCZ). The regional role
of the convergence is well seen in the southern Cook Islands: in winter
(May-October) the SPCZ generally lies to the north of the group and dry,
trade-wind dominated conditions prevail whereas in summer (NovemberApril) the convergence generally lies further south over the islands, producing
unsettled weather, heavy rainfall and, in some years, destructive tropical
cyclone activity. Typically, therefore, 6 1-68"/0 of the annual rainfall falls in
the summer months in the southern Cook Islands, and 32-390/0 in the winter
(Thompson, 1986). Further east, the Pitcairn Group generally does not show
such seasonal variability in rainfall (Fig. 9): for the period 1940-1971, winter
(May-October) rainfall averaged 903.5 mm and summer (November-April)
894.6 mm, an almost exact 50 : 50 split (Taylor, 1973). Nevertheless, several
lines of inquiry suggest that migrations of the SPCZ, and its relation to the
SOI, are significant in this area. As the SPCZ is an area of convergence, it
is associated with an area of high cloudiness (Trenberth, 1976) whose
variability can be assessed from the patterning of bright bands on daily
satellite images. Using such imagery, Streten (1973) was able to show that
the westerly migration of the SPCZ between 1969 and 1971 was accompanied
by lower than average rainfall at Pitcairn Island (- 18%) and increased
rainfall at Rapa (+@Yo; 6 months with rainfall >50% above normal) to the
south west. These findings were confirmed by Meehl (1987) who demonstrated
that cold phase periods at Pitcairn over the period 1940-1981 were associated
with rainfall deficits of at least 10 mm month-' and up to 20 mm month-'.
The explanation for these differences is that during a cold phase ENS0
period the strong east central Pacific high pressure system displaces the SPCZ
to the south and west. Stations to the east, therefore, experience higher sea
level pressures, lower sea surface temperatures and less rainfall. By contrast,
during a warm phase the weakening of the high pressure allows the SPCZ
to migrate north and east, bringing lower sea level pressures, higher sea
surface temperatures and increased rainfall to these areas (Fig. 10; Meehl,
1987). High pressure is displaced south and at 25-30"s blocks the low
pressure systems usual at these latitudes (Trenberth & Mo, 1985). The low
pressure systems move north of these blocking highs and their troughs extend
into the tropics, bringing anomalous northerly and westerly winds (Finney,
1988; Finney et aZ., 1989). These fluctuations have now been more clearly
demonstrated from the estimation of pentad (5-day) and monthly rainfall totals
in oceanic areas using remotely-sensed variations in cloud-top temperatures and
outgoing longwave radiation. With this methodology, Janowiak & Arkin
T. SPENCER
26
Januarv
0'
15"
30"
0"
15'
30"
Annual
0"
15"
30"
Figure 9. Summer (January), winter (July) and annual precipitation for South Pacific
stations (mm). After Taylor (1973).
(1991) have documented large-scale shifts in rainfall patterns in a transition
from warm phase (1986-1987) to cold phase (1988-1989) conditions (see
Gaffney, 1991; Lavery & Karoly, 1992 for detailed chronology). Comparison
of estimated austral summer rainfall (December to February) for the years
1986-87, 1987-88 and 1988-89 shows the progressive westward migration of
the SPCZ and the weakening of connective activity as the system moved
towards high values of the SOI. In the winter months (June-August) rainfall
totals rose in the region of the Pitcairn Group.
Finally, there is also a good correlation between the SO1 in September
and the distribution and frequency of tropical cyclones in the following
season. When the SO1 is positive (i.e. in a cold phase) cyclogenesis occurs
between 125"E and 180" whereas when the SO1 is negative (warm phase)
cyclones orignate in a broad band between 145"E and 120"W (Drosdowsky
& Woodcock, 1991). In the exceptional ENS0 warm phase of 1982-1983,
the median location for cyclogenesis was 11"s 162"W as opposed to the
median coordinates for all cyclones on record of 14"s 170"E (Revel1 &
Goulter, 1986). The isolated position of the Pitcairn Group has, until recently,
placed it beyond systematic monitoring for cyclone activity: the nearest
regularly observed area, the grid cell defined by 20-25"S, 150-155"W,
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
27
a) Strong annual cycle
b) Weak annual cycle
Figure 10. Cartoon illustrating processes during a) cool/high SO1 index phase and b) warm/low
SO1 index phase of the El Nino Southern Oscillation. After Meehl (1987).
experienced only 16 tropical cyclones in the 40 seasons between November
1939 and April 1979 (Kerr, 1976; Revell, 1981) and numbers might be
expected to be lower than this total further east. However, occasional minor
cyclone activity has been reported from the Pitcairn region (e.g. cyclone
Hinano in February 1989) and intense storms and hurricanes from lower
latitudes decay in this area (e.g. Hurricane Peni, February 1990) (Ready &
Woodcock, 1992).
ENS0 episodes and rainfall at Pitcairn Island
A study of the monthly rainfall totals for Pitcairn Island allows the
hypothesis of ENSO controls on the climate of the Group to be investigated
T. SPENCER
28
"
c
.t"40
1111111
43
46
49
52
55
58
61
67
Year
64
70
73
76
Figure 11. Mean annual rainfall at Pitcairn Island. Long-term mean
=
79
82
85
88
91
1716 mm.
TAIH.I:.
1. Dry periods (consecutive monthly total below 100 mm) at Pitcairn Island, 19401992
No of
months
14
11
7
0
Period
Rainfall total
(mm)
June 197ii-July 1976
March 1978-January 197!1
September 1976-March 1977
June 1970-November 1970
61 1
472
3!)2
2ox
~-
~
Expected rainfall
total* (mm)
~~
2035
1,574
986
X6 1
Vo difference
actual-expected
~~
- 70
- 70
~
60
- 70
~~
*Total rainfall that might have been expected over the drought period, assuming monthly total in
each month was long-term monthly mean (1940-19!)2).
in more detail. Rainfall records are available from March 1940 until July
1992, although monthly totals are largely missing between February 1946
and December 1953, and from part of 1986. The long-term (1941-45, 195485 and 1987-91) mean annual rainfall at Pitcairn is 1716 mm but with
considerable inter-annual variability (Fig. 11): the totals for the highest 3
years on record are 2630 (1958), 2580 (1984) and 2296 mm (1973) compared
to 942 (1977), 669 (1976) and 608 mm (1978) for the 3 years with the
lowest totals. The variation in monthly total ranges from 589 mm (October
1980) to 3.5 mm (February 1983). Defining a dry month as one with less
than 100 mm of rainfall identifies several significant drought periods (Table
1). It is worth noting that were it not for the month of August 1976, a dry
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
29
h
ri
Q,
2I
ri
4
40 1
::
a" - 8 0 - 1 ' "
40
43
1 1 I
46
I
49
I
I
/
'
55
58
61
~
~ 1 1 1~ 1 1 1~
~~~~~~
~
~ 1 1 ' 1 1 1~ 1 1 ~1 1 1 ~1 1 1 ' '
64 67 70 73 76 79 82 85 88 9 1
Year
Figure 12. Percentage deviations of individual annual rainfall totals from long-term mean,
Pitcairn Island.
52
~
period would have stretched for 22 months from June 1975 to March 1977;
and in the 26 months to July 1977 only two monthly totals exceeded 100
mm. This anomaly is clearly seen by plotting the percent deviations for
individual years from the long-term mean rainfall total (Fig. 12).
To see if these variations might be related to ENSO events requires some
re-casting of the rainfall dataset, and in a way that allows for an incomplete
record. Ropelewski & Halpert's (1987) analysis of monthly precipitation
records from the Pacific basin has identified a series of core regions in which
it is possible to identify the period within an ENSO cycle which exhibits
the largest signal and the sign of that signal. Their analysis (1) confirms the
broad pattern described above, in that for the nearest core region to the
Pitcairn Group, the South Central Pacific region, the warm phase signal is
associated with increased precipitation and (2) shows that the signal is clearest
in the period from the start of the warm episode (July) to the following
June (Fig. 13), probably in part a result of the lagged northerly movement
of the SPCZ (Rasmussen & Carpenter, 1982). Thus the Pitcairn precipitation
years were redefined to run from July to the following June. The data in
each year so defined were then ranked from 1 (the year with the lowest
rainfall) to n (the year of maximum precipitation). Each year was then
assigned a precipitation percentile index where the index for the year with
rank m was 100 m/n.(Fig. 14). When strong and medium warm phase ENSO
episodes from Quinn's chronology (1992) are identified on this plot then it
~
T. SPENCER
30
J A S O N D J F M A M J J A S O N D J FMAM J
(-1
(+I
(0)
Figure 13. Precipitation index from South Central Pacific stations, indicating higher monthly
rainfall totals in year following onset of warm phase ENSO conditions. After Ropelewski &
Halpert (1987).
is clear that there is a general correspondence between warm phase ‘years’
and wetter conditions. Also, with one exception (1956) strong cold phase
years (1955, 1964, 1970, 1971, 1973, 1975, 1988: when the SO1 index
remained in the upper 25% of its distribution for 5 months or longer
(Ropelewski & Jones, 1987)) were associated with lower than average rainfall
totals. Nevertheless, there are warm phase years which are anomalously dry
(of which 1976 is the most striking example) so ENSO linkages can only be
a partial answer to explaining the variability of Pitcairn Island’s rainfall.
It is unfortunate that this record cannot be evaluated in a longer-term
context: it is interesting to speculate as to whether rainfall totals at Pitcairn
followed those of Rarotonga and Apia which both show substantial inter-
Large scale ENSO events (after Quinn. 1992)
1940
1950
1960
1970
1980
1990
Figure 14. Rainfall percentile index for Pitcairn Island rainfall and warm phase ENSO events.
For methodology and discussion see text.
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
31
decadal variations with dry conditions in the late nineteenth and early
twentieth centuries and high rainfall totals in the 1920s and 1930s (Stoddart
& Walsh, 1992). Similar percentage changes to those experienced at Rarotonga
for Pitcairn would yield long-term mean annual rainfall totals ranging between
1980 mm (+15.4%) and 1590 mm (-7.4%).
Palaeoclimates of the Pitcairn Islands: some speculations
There has been considerable speculation as to the nature of past climates
in the south-central and south-east Pacific, particularly with reference to
conditions for Polynesian voyaging and island colonization. Particular interest
has focused on the climatic fluctuations of the Little Climatic Optimum
(LCO; 750-1250 AD), when global mean temperatures were 0.5 to 1.0"C
higher than present, and the following Little Ice Age (LIA; 1400-1850 AD)
when temperatures were c. 1.5"C cooler than during the preceding warm
period (Grove, 1988). O n the basis of rather limited evidence, Bridgman
(1983) has speculated that the LCO was characterized by a poleward shift
in the subtropical high pressure belt and a corresponding poleward shift in
the doldrums, leading to 'fair weather' westerlies, clear skies and few storms
and thus ideal conditions for long-distance voyaging. Conversely, voyaging
ceased during the LIA because of increased variability in weather systems,
including erratic trade winds and increased storminess. Finney ( 1Y85), however,
has argued that the main phase of colonization preceded the LCO, with
voyaging taking advantage of westerly winds during warm phase E N S 0
conditions; he finds no evidence for a change in the frequency of warm
events during the period of Polynesian colonization of the Pacific. Furthermore,
recent theories that the colonization process was facilitated by contrary trade
winds on outward voyages (Irwin, 1993) suggests that stronger trade winds
might not be such a hindrance to the colonization process as has been
supposed. This debate clearly needs to be resolved with improved chronologies
of Pacific island settlement, and more comprehensive reconstructions of
environmental histories for these islands; however, separation of environmental
degradation as a result of climate change from the consequences of ecosystem
manipulation by Polynesian colonists is likely to be a difficult task.
O n much longer timescales, reconstruction of sea surface temperatures for
the Last Interglacial ocean (isotope stage 5e; c. 125 kyr BP) suggests little
change from present conditions: 60% of the palaeotemperature estimates
gained from microfauna within deep-sea core sediments differ from present
day values by less than the typical
1.0 to l.,5"C standard error of estimate
(CLIMAP Project Members, 1984). However, palaeotemperature estimates
obtained from core V19-53 (17"s 113"W) suggest a southern summer
(February) mean sea surface temperature of 293°C and a southern winter
(August) temperature of 27"C, in each case 39°C warmer than the equivalent
present sea surface temperature at this latitude. Periods of higher sea surface
temperatures at the present time in this region lead to higher rainfall (see
above). However, not all the deep-sea core evidence suggests such changes
in temperature (Table 2).
For interstadial events after the Last Interglacial and at the last glacial
maximum analyses of deep-sea core materials and fossil molluscs suggest that
32
T. SPENCER
TABLE
2. Sea surface temperatures in selected mid-latitude deep-sea cores: isopotic-stage 5e
(Last Interglacial) compared to modern estimates
Core
Location
.
~~
V 1%.53
(Moore et al.,
1980)
v32-126
(Thompson,
1981)
V2 1- 146
(Moore et al.,
1980)
SST,, ("C)'
Atlas'
Core Top'
5e'
SST,("C)'
Atlas'
Core Top'
5e'
.
~~
17"Ol'S
113"31'W
27.0+3.0
23.1
23.9
29.3 & 1.5
25.4
28.3
35"19'N
117"55'E
18.9k 1.Fi
16.0
16.8
27.0+ 1.5
24.6
25.2
37"41'N
163"02'E
17.3k2.4
14.5
13.5
24.X+ 1.8
24.0
22.0
_ _ _ _ ~
SST, = winter sea surface temperature in February (N. Hemisphere) or August (S. Hemisphere).
SST, = summer sea surface temperature in August (N. Hemisphere) or February (S. Hemisphere).
Last Interglacial seasonal sea surface temperature (with standard error of estimate) from transfer
functions applied to deep sea cores.
'Modern SST from atlas values.
Holocene SST from core top sediments.
I
the Southern Ocean as a whole was c. 2°C cooler, although it seems that
little temperature change affected the mid-latitude gyre of the south-central
Pacific (CLIMAP Project Members 1976; and see Spencer, 1989 for review).
Rather, the northern hemisphere glaciations led to intensification of the
atmospheric-oceanic circulation systems: studies of quartz abundance
distribution in core V19-29 (3"S, 83'56'W) have suggested that the trade
winds were more intense during the glacial phases, particularly between 7361 kyr BP and 43-16 kyr BP (Molina-Cruz, 1977). Changing distributions of
radiolarian assemblages in deep-sea cores indicate that the strengthening of
the glacial trades was accompanied by a more meridional pattern of wind
stress and an intensification of the equatorial surface circulation; in particular
cool waters associated with upwelling eastern boundary currents showed
increased westward penetration, reaching their greatest extension at the glacial
maximum (Moore et ah, 1981; Rornine & Moore, 1981). Reconstructions of
the differences in sea surface temperatures between present and glacial
maximum conditions, therefore, suggest changes were concentrated in sensitive
ocean margin locations with little change at mid-ocean sites (Spencer, 1989).
ACKNOWLEDGEMENTS
The Sir Peter Scott Commemorative Expedition was generously supported
by the following major sponsors: The Royal Society, International Council
for Bird Preservation, British Ornithologists' Union, J. A. Shirley, Foreign &
Commonwealth Office UK, UNESCO; all other sponsors are acknowledged
in the Expedition report of 1992. This is paper 30 of the Sir Peter Scott
Commemorative Expedition to the Pitcairn Islands.
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APPENDIX A: MAPS OF THE PITCAIRN ISLANDS
i
G
contour interval
- 50 metres
kilometre
0
I
~
~
Figure A l . Pitcairn Island.
I
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
I
21
I
128'20'W
I
19'
37
I
1s'
19
24'205
21
22
Figure A2. Henderson Island. Island outline and topography from Spencer & Paulay (1989)
and SPOT satellite imagery. Vegetation communities and mapping by S. Waldren; cartography
by S. Waldren, J. Wyatt and T. Spencer.
3x
T. SPENCER
130'45 W
130"44
Boat
/-y
I
.
23'51
1
S w a n a beachfront
Sand flats
Sand spit
Shoreline scrub
63
cz1
Inland Argusia scrub
0
Tall Argus/a forest
Rson/a/Argus/a forest
Figure A3. Oeno Atoll. Reef and land area positions from Spencer ( I W ~ ) ,GPS surveys by
Expedition personnel (A. Hendricsen, S. Schubel) and aerial photography. Vegetation
communities by S. Waldren. Cartography by S. Waldren, M. Critchley (EFWMaptec Ltd.),
J. Wyatt and T. Spencer.
-
-
Lagoon
full of
coral Ridges
/
7
I
Figure A4. Ducie Atoll. Reef position from Admiralty Chart; land areas from GPS surveys
by Expedition personnel (A. Hendricsen, S. Waldren). Cartography by S. Waldren,
M. Critchley (EKA/Maptec Ltd.), J. Wyatt and T. Spencer.
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GKOUP
39
APPENDIX B: THE CLIMATE O F HENDERSON ISLAND 199-1992
The Expedition established a standard meteorologcal station at 3 0 m above present sea level above
the North Beach of Henderson Island. Daily measurements were taken at 1x00 UT (0900 local time).
One year's meteorologcal data were collected, from 1 February 1991 to 31 January 1992. The following
datasets and observations provide background information o n the climate of Henderson; its relation to
the climate of Pitcairn Island; and the character of the Expedition's 'weather year' in the context of
wider climatic oscillations within the Pacific Ocean basin.
Winindspeedc and direction
Measurements of wind speed and direction were made on 320 days (88%~of the year). At time of
recording, windspeeds in excess of ,5 m s ' were recorded on 102 days (32%~of the record) and in
excess of 10 m sCI on 24 days (7..5V11).Calms were rare. There was no particular pattern to the
distribution of these windspeeds by month. Winds from an easterly direction dominated the record
(4(Nu of the time) but winds from all directions were experienced with a considerable proportion of
northerly (21%1)and southerly (13%~)
winds. This may have been due to warm phase ENSO conditions
(see below).
These observations can be compared with a 10 year record from Pitcairn Island (197!)-86+ 1!)8990, a = 31 19 daily observations); this confirms the findings of the Henderson record but in addition
allows some disaggregation of a more comprehensive dataset. Northeasterly and easterly winds are
more prevalent during the austral summer from November to April whereas the winter months (MayOctober) are characterized by a much greater variety of wind directions (Table B l ) . Winds exceeding
5 m sCI are typically experienced on 1\5-16 days per month, with no discernible seasonal variation.
in 1979; 142 days in 1!)00).
Inter-annual variability is, however, marked (e.g. 212 days > 5 m s
Calms are rare (2%~of days) and gale force winds ( > 17.2 m S C ' ) even more so, with only 17 days
of gales, largely in the months of May, June and July, over the entire 10 year period (O.S1,/lof days).
Van Loon & Shea (1987) have argued that the interaction between anomalous sea surface temperatures
and atmospheric circulation systems in the spring (November-January) preceding ENSO warm phases
produces westerly winds. This appears to have been the case in 1990 and in 1986 when westerlies
accounted for 21.4"h and 24.8'%1 of wind directions respectively, compared to the 10-year mean of
17.7"!.
'
Temperature
Temperatures showed a clear seasonal pattern; the mean monthly maximum temperature varied
between 29.6"C (February 1991) and 242°C (June 199 1); the comparable minimum temperatures were
222°C (February) and 15.7"C (June) (Fig. Bl). The maximum temperature recorded was 31.4"C, on
three occasions in February and March, and the absolute minimum 12.0"C (September). On Pitcairn
Island for the corresponding period the mean maximum temperature range was 25.1"C (February) to
10.4"C (June) and the mean minimum temperature ranged between 20.6"C (March) and 16.l"C (June).
Mean monthly maximum temperatures were consistently higher on Henderson Island (x = 4.7"C;
range = 4.2-5.1"C); mean monthly minimum temperatures were also higher than those on Pitcairn,
T~\EI,I:
B 1. Percentage frequency of wind directions, Pitcairn
Island, 1979-86+ 1989-90
Direction
(degrees)
November-April
340-029
030-069
070-1 10
120-159
160-209
210-249
2.50-299
300-339
Calms
May-October
Complete year
~..
~~~
20.7
25.0
20.5
12.4
8.X
2.7
3.3
4.6
2.1
23.3
13.7
13.0
11.x
11.x
10.5
6.3
7.7
2.0
23.0
19.2
16.7
12.1
10.3
6.7
4.8
6.2
2.0
T. SPENCER
40
Q,
L
3
30
4-
a
L
0)
n 25
E
20
15
10
Figure B1. Mean monthly maximum (open squares) and mean monthly minimum (closed
circles) temperatures, Henderson Island, February 1991-January 1992.
although less so ( x = 1.1"C; range = -0.4"C to +2.1°C). These differences may be due in part to the
different heights of the recording stations (Henderson -30 m a d . ; Pitcairn -264 m a.s.1.). In the
south-west Pacific, the environmental lapse rate is 8.9"C in January and Y . Y T in July (Tomlinson,
1975); these figures suggest that about half the observed difference might be explained by this control.
When compared with the long-term record (1940-45+ 19.56-86+ 1989-90), the mean annual temperature
at Pitcairn Island for the year February 1991-January 1992 was 0.5"C cooler than expected (x-20.7"C;
long-term x = 21.2"C). In particular, the low mean monthly and mean monthly minimum temperatures
for June and July 1991 and January 1992 were outside one standard deviation of the comparable longterm means. It has been suggested (T. Benton. personal communication 1992) that this may have been
due to warm phase E N S 0 conditions in the Pitcairn region at this time '(see below), a hypothesis
worthy of further study.
"V
Henderson
Pitcairn
200
100
0
Figure B2. Monthly rainfall at Pitcairn Island and Henderson Island, February 1991-January
1992.
GEOLOGICAL AND CLIMATIC BACKGROUND, PITCAIRN GROUP
41
T.zei.i. B2. Duration of wet ( > 5 mm day-')
and dry ( <5 mm day-') spells at Henderson
Island, February 1991-January 1992
Duration of spell
(days)
~-
Number of occurrences
Wet
Dry
-
1
2
3
4
5
6
7
8
1)
10
11
12
13
14
15
lh
17
18
~
31
12
2
8
7
h
5
2
2
1
3
0
2
3
0
0
0
1
0
0
2
Prec$itation
The rainfall total for the period February 1991 to January 1902 at Henderson Island was 1623 mm.
In the same period Pitcairn Island recorded 2171 mm (+34%1; Fig. B2). The three highest 24-hour
totals were unexceptional: 58.1 mm (lW,5/91), 64.2 mm (1/2/9l) and 65.1 mm (2~5/12/91).66"h of the
total raindays recorded less than ,5 mm, 77"h less than 10 mm and 95"h less than 30 mm. Durations
of wet and dry spells are detailed in Table B2.
Longer-term context
In terms of climatological statistics, was the period February 1991-January 1992 an average year?
Variations in the Southern Oscillation Index between January 1987 and May 1992 show that the Index
30
I
1
so1
v 5 month mean
I
20
10
0
- 10
-20
!
Jan
-30
1987
1988
I
1989
I
1990
I
1991
I
1992
I
Jan
Jan
Jan
Jan
Jan
Figure B3. Southern Oscillation Index, 1990- 1992. Note negative values (warm ENSO phase)
whilst Expedition was in the field.
T. SPENCER
42
V
poor
I
Feh-1
1
1
Apr-1
1
1
I
I
J u n - 1 Aug-1
'
I
I
1
Oct-1 Dec-1
Month
'
1
Feh-1
'
I
Apr-1
"
Jun-1
Figure B4. Percentage deviation of monthly rainfall totals from corresponding long-term
monthly mean, Pitcairn Island, February 199 I -January 1992.
became negative (i.e. moving towards a warm phase) in February 1991 (Beard, l993), precisely when
Expedition meteorological records commenced. Positive sea surface temperature anomalies developed
in the equatorial Pacific by September and reached +2 to +3"C along the Peruvian coast by November
1991, typical of a maturing warm phase (Rasmussen & Wallace, 1983). The SO1 continued to fall
throughout the field phase of the Expedition, only recovering to a positive value in May 1992 (Fig.
B3; Wright, 1993). Clearly, therefore, the Expedition experienced a warm phase, or El Nino event,
for its duration. Typically, higher than average rainfall follows in the year after the onset of a warm
phase episode (year defined as July (0) to June(+); see analysis above). In this case there was some
evidence for this pattern in the percent deviations from long-term monthly means in the Pitcairn Island
record: positive deviations characterized September, October and December 1991 and January and
February 1992, although consistent negative anomalies persisted thereafter to June 1992 (Fig. B4).
The analysis of Pitcairn's rainfall record has shown that the mean annual rainfall can fall to as low
as -600 mm and that drier than average years can persist for long periods (e.g. 1975-1980). If
Henderson Island's precipitation is typically 30Vu less than that of Pitcairn Island then there are likely
to be times when the climate of Henderson Island must be extremely dry. This should be borne in
mind when considering the ecological studies which follow in this volume.
Acknowledgements
I am g~ateful to all Expedition members who collected meteorologcal data in the field; to Dr T.
G . Benton for preliminary analysis of meteorologcal records; to I. Hephurn for access to his laboriously
transcribed meteorological records from Pitcairn Island; to the Library, Meteorological Office UK for
assistance with tracing historical records; and to T. Christian for the supply of missing record data
from Pitcairn Island.