these en cotutelle

Transcrição

these en cotutelle

N° d’ordre: 3344
THESE EN COTUTELLE
Universidade de Lisboa
Faculdade de Ciências
Departamento de Geologia
Université Bordeaux
Ecole Doctorale des
Sciences du vivant, Géosciences
Science de l’environnement
Présentée à
L’UNIVERSITÉ de LISBONNE
Par Filipa
Naughton
POUR OBTENIR LE GRADE DE
DOCTEUR
. en Géologie spécialité Paléontologie et Stratigraphie (Portugal)
. en Océanographie, Paléo-Océanographie (France)
As variações climáticas dos últimos 30 000 anos e sua influência
na evolução dos sistemas costeiros do norte de Portugal
Les variations climatiques des derniers 30 000 ans et leur influence sur
l’évolution des systèmes côtiers du nord de Portugal
Dirigée par :Mme Maria Fernanda SÁNCHEZ GOÑi et Mme Maria da Conceição FREITAS
Soutenue le: 18 Janvier 2007
Devant la commission d’examen formée de :
Président du Jury :
M. J.-L. TURON, Directeur de Recherche CNRS (EPOC-Université Bordeaux 1)
Rapporteurs :
M. H. SEPPÄ, Maître de Conférence (Dép. de Géologie-Université Helsinki)
Mme F. ABRANTES, Directeur de Recherche (Dép. de Géologie Marine-INETI, Lisbonne)
Examinateur :
Mme T. DRAGO, Chercheur Auxiliaire (IPIMAR, Olhão)
Directeurs de thèse :
Mme M.F. SÁNCHEZ GOÑi, Maître de Conférence EPHE (EPOC-Université Bordeaux1)
Mme M.C. FREITAS, Professeur Associé-FCUL (Université de Lisbonne)
Principalmente para ti....................
Maria da Luz Naughton
Agradecimentos, Remerciements
A execução do presente trabalho não teria sido possível sem a contribuição
de diversas pessoas e entidades, a quem desejo expressar os mais sinceros
agradecimentos, nomeadamente:
A Maria Fernanda Sánchez Goñi pour m’avoir contaminée avec sa passion
pour la recherche et soutenue au cours de ces derniers quatre ans. Je te remercie,
Maria, pour toute ta gentillesse et ton amitié, pour avoir toujours été disponible et
aidée à n’importe quelle heure, chaque jour de la semaine voire les week-ends.
Merci, Maria, pour m’avoir appris non seulement les pollens mais aussi beaucoup de
choses sur la paléoclimatologie. Merci, Maria, pour m’avoir fait voir le vrai travail en
équipe, avec toutes ces discussions scientifiques partagées entre autres dans la salle
à café.
À Teresa Drago a quem gostaria de expressar toda a minha gratidão e sem a
qual este trabalho não teria sido possível de ser realizado. Obrigada Teresa por teres
apostado na minha formação em palinologia, pela confiança em mim depositada,
pelo interesse e dedicação com que tens acompanhado o meu trabalho e pelo
teu encorajamento.
A Stéphanie Desprat, pour m’avoir toujours encouragée et principalement
pour m’avoir soutenue au cours de ces 4 ans. Je te remercie « Stephy » pour avoir
parfois été ma tête et pour avoir retrouver tous mes fichiers perdus dans l’immensité
de mes dossiers sur l’ordinateur. Je te remercie pour le temps passé à m’apprendre
les logiciels compliqués et à partager tes connaissances et ton enthousiasme pour la
recherche. Surtout, je te remercie pour cette complicité et notre amitié, nos
discussions parfois « trop » scientifiques dans les cafés de la place Camille Julian …
heureusement qu’on ne peut pas discuter dans la salle du cinéma UTOPIA …
À Professora Conceição Freitas pela sua disponibilidade e co-orientação,
pelas suas críticas e sugestões e pela sua simpatia sempre presente.
A Josette Duprat, pour avoir été toujours là pour partager ses résultats
magnifiques de foraminifères planctoniques ainsi toute sa connaissance sur les
changements des ces associations. Merci Josette pour votre dynamisme et votre
sens pratique afin de répondre à toutes les questions scientifiques posées sur le
moment. Je vous remercie aussi pour toute votre gentillesse …
A Marie-Hélène Castera, pour les préparations des nombreuses lames
palynologiques, son dynamisme et sa bonne humeur. Je te remercie pour le temps
que tu as passé sur mes lames incluant le montage d’une deuxième, troisième voir
sixième lame sur un même niveau. Il faut le dire : « quel courage ! ». Pour la dernière
année, ta nouvelle partenaire de préparations d’échantillons, Muriel, a
heureusement beaucoup participé à finaliser quelques séries d’échantillons pour
moi et je la remercie aussi vivement.
A Jean-Marie Jouanneau, pour son soutien et son assistance quotidiens. J’ai
beaucoup apprécié votre intérêt pour mon travail et je vous remercie pour toutes les
discussions et les échanges sur de nombreux sujets scientifiques, vous avez contribué
par vos relectures et critiques à l’aboutissement des articles scientifiques.
A Jean-Louis Turon pour m’avoir reçu en toute gentillesse au sein de votre
équipe « Biopal », pour m’avoir donné la possibilité de travailler sur des carottes
Bordelaises et pour m’avoir passé votre magnifique microscope. Je vous remercie
aussi pour les donnés sur les assemblages de dinokyste de la carotte Française.
A l’équipe des traceurs isotopiques, en particulier Karine Charlier et Bruno
Malaisé pour les donnés isotopiques. A Francis Grousset pour les discussions sur les
événements d’Heinrich et son étudiante Elsa Jullien pour avoir toujours été disponible
pour partager ses connaissances acquises pendant cette dernière année et pour les
discussions de « paléoclimapoésie ». Je tiens aussi à remercier aussi Philippe Martinez
pour m’avoir proposé une vacation qui a contribué à financer partiellement ma
dernière année de thèse.
A l’équipe d’environnement côtiers, en particulier à Monsieur Castaing pour
m’avoir aidé sur le sujet des estuaires et pour m’avoir permis de participer à
quelques sorties de terrains ainsi qu’à quelques cours de DEA. Je tiens aussi à
remercier Sébastien Zaragosi pour avoir toujours été disponible à m’aider et à
discuter sur les images radiographiques des carottes marines et aussi sur les données
obtenues de l’étude sédimentologique des lames minces. A Michel Cremer pour la
gentillesse de m’aider sur les images radiographiques des carottes marines.
Je tiens aussi à remercier tout le couloir BIOPAL pour ces dernières années
passées dans la bonne humeur et pour m’avoir reçue et intégrée au sein de leur
équipe.
A Jean-François Bourillet pour m’avoir donné la possibilité de travailler sur une
de ses magnifiques carottes (IFREMER-Brest), dans le cadre d’un contrat de travail de
3 mois lequel à contribuer partiellement à financer ma dernière année de thèse. Je
vous remercie aussi pour m’avoir intégré dans l’équipe d’étude de cette carotte et
de m’avoir permis l’écriture d’un article scientifique.
A Elsa Cortijo et Elisabeth Michel du Laboratoire des Sciences du Climat et de
l’Environnement (LSCE), Gif-sur-Yvette, pour les analyses isotopiques et pour toujours
avoir été disponibles à résoudre des questions scientifiques du moment.
A Frauke Rostek et Edouard Bard (Cerege) pour les données des alkénones.
A Vincent Marieu pour avoir perdu pas mal de temps à m’apprendre
l’analyse spectrale et d’autres choses comme corriger mes fautes d’orthographe,
etc. Je te remercie aussi pour toute ta gentillesse et ton amitié. Et je m’excuse
d’avoir fumée quelques cigarettes dans notre bureau.
A Delphine Denis et Elodie Marchès pour avoir toujours été disponibles pour
m’aider à corriger mon français catastrophique ainsi que pour m’avoir montré qu’il
existe encore des étudiants passionnés et dédiés à la recherche tout en restant
simples. Je vous remercie aussi pour votre amitié et pour m’avoir supporté dans les
moments plus difficiles. Je te remercie aussi Delphine pour m’avoir beaucoup aidé
les derniers jours de cauchemar………..et pour la nuit blanche.
A Anne-Laure Danieu pour avoir aussi toujours été disponibles pour m’aider à
corriger mon français principalement dans les derniers jours de cauchemar…. Je te
remercie aussi pour les conversations partagées à la salle à café et pour les
moments d’amitié partagés.
A William Fletcher for helping me to improve my catastrophic and terrible
English.
Je tiens aussi à remercier mes deux autres collègues de Bureau : Marc,
Vincent et Jonathan pour m’avoir supporté pendant quelques années et aussi pour
leur amitié.
Je tiens aussi à remercier Alexandra Coynel pour m’avoir écouté et partagé
quelques moments très agréables.
À Pauliana Valente por ter passado algumas horas a corrigir o meu português
de actual emigrante em França.
À Àurea Narciso pela sua disponibilidade e ajuda em relação aos problemas
administrativos relacionados com a tese.
Tenho ainda necessidade de agradecer a uma série de pessoas que me
ajudaram, apoiaram há cerca de 4 anos (durante o meu mestrado) as quais
contribuíram em muito para a decisão de iniciar este trabalho de Doutoramento,
nomeadamente:
A Luisa Santos da Universidade da Coruña, aos colegas do grupo DISEPLA:
Rui Taborda, à Anabela Oliveira e João Cascalho e Francisco Fatela.
Gostaria ainda de agradecer:
Au UMR CNRS 5805 EPOC (Environnements et Paléoenvironnements
Océaniques- Univérsité Bordeaux 1) pour m’avoir reçue pendant 4 ans et de m’avoir
donné toutes les conditions favorables à la réalisation de mon travail de thèse.
À Universidade de Lisboa por ter aceite a minha candidatura a
Doutoramento em regime de co-tutela.
Ao IPIMAR- Instituto de Investigação das Pescas e do Mar por me ter acolhido
desde o meu mestrado e início do meu Doutoramento.
Gostaria de expressar ainda a minha maior gratidão às entidades e projectos
que financiaram este trabalho de Doutoramento, nomeadamente:
O suporte financeiro durante 3 anos do projecto -“Envi-Changes”. Late
Quaternary Environmental Changes From Estuary and Shelf Sedimentary Record
inserido no Programa Dinamizador das Ciências e Tecnologias do Mar
(PDCTM/PP/MAR/15251/99) e no Programa de Apoio à Reforma dos Laboratórios de
Estado da Fundação para a Ciência e a tecnologia (FCT), o qual foi coordenado
pela Doutora Teresa Drago.
À cooperação Luso-Françesa (ICCTI- IFREMER) por me ter financiado algumas
missões entre Portugal e França.
Ao programa “PESSOA” integrado no programa, de acções integradas
franco-português por me ter financiado algumas missões entre Portugal e França.
À embaixada de França em Portugal por me ter apoiado financeiramente 2
meses no início do meu 4° ano de Doutoramento.
O projecto CNRS-ECLIPSE “La variabilité climatique d’ordre millénaire de la
dernière période glaciaire dans les moyennes latitudes de l’Atlantique Nord, d’après
l’analyse pollinique d’une carotte marine profonde prélevée sur la marge ibérique »
o qual me proporcionou uma « vacation » de um mês.
O programa EuroCLIMATE-European Science Foundation.
L’Agence National de la Recherche (ANR)
Finalmente gostaria ainda de agradeçer:
À minha avó Maria da Luz Naughton pelo seu apoio incondicional e pela
força que toda a vida me deu.
E a:
Heikki Seppä and Fátima Abrantes for their presence as the principal jury
members during my PhD defense and for their valuable comments which greatly
improved this manuscript.
Resumo
O principal objectivo deste trabalho é caracterizar a variabilidade climática
que ocorreu nos últimos 30 000 anos, nas médias latitudes do Atlântico Norte. Este
trabalho focaliza-se ainda no impacto da variabilidade climática Holocénica na
evolução dos sistemas costeiros do noroeste de Portugal. Desta forma, foi
efectuado um estudo multidisciplinar (ex: pólens, associações de foraminíferos
planctónicos, δ18O de foraminíferos planctónicos e bentónicos e alcanonas) em
duas sondagens marinhas profundas recolhidas no noroeste da margem Ibérica. Foi
ainda efectuado um estudo “multi-proxy” em duas sondagens estuarinas do
noroeste de Portugal assim como numa sondagem marinha pouco profunda
recolhida na plataforma continental noroeste Francesa. Foi igualmente efectuada
uma calibração do sinal polínico ao longo da margem Ibérica.
Este estudo mostra que os eventos associados a fortes descargas de
icebergues no Atlântico Norte, designados por eventos de Heinrich, são complexos
e compostos essencialmente por duas fases climáticas distintas no noroeste da
margem Ibérica e continente adjacente. A primeira fase é marcada por um
arrefecimento extremo das condições atmosféricas no continente (declínio da
floresta de Pinus) e oceânicas de superfície. Esta fase é ainda caracterizada no
continente, por uma certa humidade a qual é representada pelo desenvolvimento
de Ericaceae incluindo Calluna e pelo aumento da concentração polínica total. A
segunda fase é marcada por condições menos frias tanto no oceano como no
continente assim como por um aumento da aridez continental (desenvolvimento de
plantas
semi-desérticas).
Foram
propostos
dois
mecanismos,
um
oceânico
(circulação oceânica) e um outro atmosférico (relacionado com a Oscilação Norte
Atlântica), para explicar o tal padrão complexo deixado pelos eventos de Heinrich
no noroeste da margem ibérica.
O último máximo glaciar apresenta um sinal oposto entre as condições
oceânicas de superfície (temperaturas relativamente elevadas) e a vegetação
continental
a
qual
foi
dominada
por
formações
abertas
(temperaturas
relativamente baixas). Este sinal oposto é explicado, por um lado, pela
intensificação da “Meridional overturning circulation” (MOC) no Atlântico Norte
(MOC), e por outro, pelo aumento do albedo, pelo forte contraste sazonal existente
nas altas latitudes do Atlântico Norte, e pela diminuição da concentração de CO2
na atmosfera. A intensificação da MOC facilitou o aumento da temperatura
oceânica, enquanto que os outros mecanismos propostos contribuíram para a
manutenção das temperaturas frias no continente.
Após o evento H1, o aumento da insolação de verão nas médias latitudes do
Hemisfério Norte e a intensificação da MOC provocaram um aquecimento
oceânico e continental (expansão da floresta decídua) o qual caracteriza o evento
Bölling-Alleröd no noroeste da Península Ibérica. Há cerca de 14 000 anos cal BP,
ocorreu um incidente extremamente quente o qual é marcado pela máxima
expansão da floresta decídua (Quercus). Este episódio é contemporâneo do
importante evento relacionado com uma subida rápida do nível do mar (Meltwater
pulse 1A) e do pico máximo observado na temperatura da Gronelândia durante o
Bölling-Alleröd. Isto sugere que este aquecimento extremo do Hemisfério Norte terá
sido provavelmente o impulsionador da fusão drástica da calote glaciar canadiana
“Laurentide ice sheet”.
A diminuição da floresta decídua de Quercus e a expansão de plantas semidesérticas caracterizam o evento frio e seco, Dryas recente, no noroeste da
Península Ibérica. Estas condições foram favorecidas pela redução da intensidade
da MOC e pelo predomínio de condições semelhantes ao modo positivo da NAOlike.
O início do Holocénico (11 600-8 200 anos cal BP) é marcado por um
aquecimento máximo no noroeste da margem Ibérica e continente adjacente e
sugere um episódio máximo térmico (HTM) nesta região.
Após o HTM, a diminuição gradual da floresta temperada sugere um
arrefecimento progressivo o qual segue o padrão geral de diminuição da insolação
de verão das médias latitudes do Hemisfério Norte.
Superimposta
a
esta
variabilidade
climática
orbital,
um
episódio
caracterizado pela expansão de Corylus indica um aumento do contraste sazonal
no noroeste de França e marca o início do evento multi-secular “8.6-8.0 ka”. No
interior deste evento, o súbito declínio da floresta de Corylus marca o evento frio
designado “8.2 Ka”. O forte contraste sazonal resulta dos episódios terminais de
expulsão dos lagos de “Agassiz” e de “Ojibway” e da consequente redução
gradual da MOC, enquanto que o súbito arrefecimento reflecte o momento de
máxima diminuição na intensidade da MOC a qual terá provocado um
arrefecimento suplementar da temperatura na Europa e na Gronelândia.
A melhoria climática característica do início do Holocénico permitiu a
expansão da floresta de Quercus na bacia do Douro e favoreceu a subida gradual
do nível do mar. Após 6 535 anos cal BP, o desenvolvimento de uma barreira
cascalhenta na parte sul do estuário resultou da atenuação da subida do nível do
mar e do aumento do hidrodinamismo do rio. A presença desta barreira contribuiu
para a migração do canal principal do rio para norte e para o alargamento deste
estuário.
Résumé
Le principal objectif de cette thèse est de caractériser la variabilité climatique
des moyennes latitudes de l’Atlantique Nord qui a eu lieu durant les derniers 30 000
ans. Cette thèse traite aussi de l’impact de la variabilité climatique sur l’évolution
des systèmes côtiers au cours de l’Holocène. Pour cela, une étude multiproxy
(pollen,
assemblages
de
foraminifères
planctoniques,
δ18O
benthique
et
planctonique, alcénones) à été réalisée à partir de deux carottes marines profondes
et une estuarienne prélevées dans le nord-ouest de la marge Ibérique et d’une
carotte marine de la plateforme continentale nord-ouest Française. Egalement, une
calibration du signal pollinique des sédiments de la marge occidentale ibérique a
été préalablement effectuée.
Cette étude montre que les événements associés à la forte décharge
d’icebergs dans l’Atlantique nord, événements d’Heinrich, sont complexes,
composés de deux phases climatiques distinctes dans la marge Ibérique et sur le
continent adjacent. La première phase est marquée par des températures marines
et continentales (chute de la forêt de Pinus) particulièrement froides et une certaine
humidité (développement des Ericaceae et augmentation de la concentration
pollinique); la deuxième, par des conditions moins froides mais une sécheresse
accrue (fréquences maximales des plantes semi-désertiques). Des mécanismes de
forçages marins (circulation océanique) et atmosphériques (l’oscillation Nord
Atlantique) ont été proposés pour expliquer ce scénario.
Le Dernier Maximum Glaciaire montre un découplage entre les températures
des eaux de surface, qui sont relativement élevées dans cette région, et la réponse
de la végétation qui est dominée par les formations ouvertes. Ce découplage est
expliqué d’une part par l’intensification de la circulation méridienne Atlantique de
renversement (MOC) observée par des études précédentes et, d’autre part, par
l’augmentation de l’albédo, le fort contraste saisonnier et la chute de la
concentration de CO2 atmosphérique, ces derniers permettant le maintien de
températures froides sur le continent.
L’augmentation de l’insolation d’été des moyennes latitudes de l’Hémisphère
Nord et l’intensification de la MOC ont produit le réchauffement climatique du
Bölling-Alleröd. Un incident extrêmement chaud dans les moyennes et hautes
latitudes de l’Atlantique Nord vers 14 000 cal ans BP est synchrone d’un épisode
important de montée du niveau marin global (Meltwater Pulse 1A). Cela suggère
que ce réchauffement de l’Atlantique Nord pourrait être le principal responsable de
la fonte drastique et soudaine de la calotte glaciaire de Laurentide.
Ensuite, l’événement froid du Dryas récent, associé paradoxalement au
maximum d’insolation estivale dans l’Hémisphère Nord, serait le résultat d’une
réduction de l’intensité de la MOC et de la prédominance du mode positif de la
NAO-like.
Le début de l’Holocène est marqué par un fort réchauffement climatique
jusqu’à 8 200 cal ans BP dans la marge Ibérique. Cette période caractérise le
Maximum Thermique de l’Holocène (HTM) sur cette région.
Après le HTM, un refroidissement atmosphérique à long-terme suit la
diminution graduel de l’insolation d’été des moyennes latitudes.
Surimposé à cette variabilité climatique orbitale, un épisode caractérisé par le
développement de la forêt de Corylus indiquant un fort contraste saisonnier dans le
nord-ouest de la France marque l’événement global pluriséculaire "8.6-8.0 ka". A
l’intérieur de cet événement, la chute soudaine de cette forêt, marque le
refroidissement du "8.2 ka". Le fort contraste saisonnier serait le résultat des épisodes
terminaux de purges des lacs d’Agassiz et d’Ojibway et de la réduction graduelle de
la MOC tandis que le refroidissement brusque serait lié à la phase ultime de
réduction de la MOC provoquant une diminution supplémentaire des températures
sur l’Europe.
Le réchauffement climatique du début de l’Holocène a permis le
développement d’une chênaie caducifoliée dans le bassin du Douro au même
temps que nous avons détecté l’augmentation graduelle du niveau marin dans son
estuaire. Vers 6535 ans cal BP, le développement d’une barrière de gravier dans la
partie sud de celui-ci serait le résultat de l’atténuation de la montée du niveau marin
et l’augmentation de l’hydrodynamisme de la rivière. La présence de cette barrière
a contribué à la migration vers le nord du chenal principal de la rivière, identifié
précédemment par des études sismiques, et par conséquence, à l’élargissement de
cet estuaire. Nous avons pu donner une date maximale pour cette migration qui a
eu lieu après 6535 ans cal BP.
Agradecimentos
Resumo
Résumé
ÍNDICE GERAL
Índice de Figuras
Índice de Tabelas
Índice de Anexos
Capítulo 1 | Introdução
1
Introduction
6
1. 1 Paleoclimatologia do quaternário recente, contexto e objectivos principais deste trabalho
12
1. 1. 1 Variabilidade climática orbital
12
1. 1. 2 Variabilidade climática milenar
15
1. 1. 2. 1 Os eventos de Heinrich e os eventos de D-O
15
1. 1. 2. 2 O início da deglaciação
31
1. 1. 2. 3 Holocénico
36
1. 2 Calibração da assinatura polínica marinha ao longo da Península Ibérica
40
1. 3 Impacto da variabilidade climática na evolução dos sistemas costeiros
44
1. 3. 1 Variações do nível médio do mar
44
1. 3. 2 Evolução dos sistemas costeiros
45
1. 4 Zona de estudo
47
1. 4. 1 Margem Ibérica
47
1. 4. 1. 1 Clima e vegetação actual
47
1. 4. 1. 2 Oceanografia
49
1. 4. 1. 3 Geomorfologia e dinâmica sedimentar actual
52
1. 4. 2 Plataforma continental noroeste Francesa
57
1. 4. 2. 1 Geomorfologia, oceanografia e sedimentação actual
57
1. 4. 2. 2 Clima e vegetação
58
1. 5 Material e Métodos
59
1. 5. 1 Amostras superficiais
59
1. 5. 2 Sondagens
62
1. 5. 2. 1 Sondagens marinhas profundas
62
1. 5. 2. 2 Sondagem marinha pouco profunda VK03-58Bis
64
1. 5. 2. 3 Sondagens estuarinas
64
1. 5. 3 Cronologia e datações
14C
1. 5. 4 Indicadores paleoclimáticos
66
68
1. 5. 4. 1 Variação do coberto vegetal e do clima continental
68
1. 5. 4. 2 Indicadores paleoclimáticos marinhos
73
1. 5. 4. 3 Variações do volume de gelo acumulado nos pólos
77
Referências
77
Capítulo 2| Present-day and past (last 25 000 years) marine pollen signal off western Iberia
105
Resumo
106
Résumé
107
Abstract
109
2. 1 Introduction
111
2. 2 Environmental Setting
112
2. 2. 1 Study area and present-day vegetation and climate
112
2. 2. 2 Oceanography
114
2. 2. 3 Morphology and recent sedimentation
115
2. 2. 3. 1 North-western Iberian margin
117
2. 2. 3. 2 South-western Iberian margin
118
2. 3 Material and methods
119
2. 3. 1 Deep-sea cores: MD99-2331 and MD03-2697
119
2. 3. 1. 1 Radiometric dating
119
2. 3. 1. 2 Marine proxy analyses
121
2. 3. 1. 3 Pollen analysis
122
2. 3. 2 Modern pollen samples
122
2. 4 Results and Discussion
124
2. 4. 1 Present day pollen signature
124
2. 4. 1. 1 Western Iberian terrestrial sites
124
2. 4. 1. 2 Western Iberian estuarine and margin sites
126
2. 4. 2 Present-day pollen transport patterns
128
2. 4. 3 Climatic and vegetational response in western Iberia to North Atlantic climatic events over the
last 25 000 years
129
2. 4. 3. 1 Marine Isotopic Stage 2
132
2. 4. 3. 1. 1 Heinrich events (H2 and H1)
132
2. 4. 3. 1. 2 The LGM
137
137
2. 4. 3. 2. 1 The Bölling-Allerød
137
2. 4. 3. 2. 2 The Younger Dryas cold event
139
2. 4. 3. 2. 3 The Holocene
139
2. 5 Conclusions
References
142
143
Capítulo 3| New insights on the impact of Heinrich events and LGM in the mid-latitudes of the eastern
North Atlantic and in the adjacent continent
153
Resumo
154
Résumé
155
Abstract
157
3. 1 Introduction
159
161
162
3. 3. 1 Chronostratigraphy
162
3. 3. 2 Pollen analysis
165
3. 3. 3 Marine proxy analysis
166
3. 3. 3. 1 Isotopic analyses
166
3. 3. 3. 2 Ice rafted detritus (IRD)
166
3. 3. 3. 3 Planktonic foraminifer-derived SST
166
3. 3. 3. 4 Alkenone-derived SST
167
3. 4 Results and discussion
168
3. 4. 1 Long-term climate variability and the LGM period
178
3. 4. 2 Heinrich events
173
3. 4. 3 Possible mechanisms triggering the complex pattern signal of Heinrich events in and off northwestern Iberia
175
3. 5 Conclusions
179
References
180
Capítulo 4| Climate variability during the last deglaciation in north-western Iberian margin and
adjacent continent
189
Resumo
190
Résumé
192
Abstract
193
4. 1 Introduction
197
198
199
4. 3. 1 Stratigraphy and age model
199
201
4. 3. 3 Marine proxy analyses
201
4. 3. 3. 1 Ice rafted detritus (IRD) and planktonic foraminiferal assemblages
201
202
4. 4 Vegetation and climate changes in north-western Iberia and adjacent margin during the last
deglaciation
203
4. 4. 1 The end of the Last Glacial maximum (LGM)
205
4. 4. 2 The Heinrich 1 (H1)
205
4. 4. 3 The Bölling-Alleröd (B-A)
208
4. 4. 4 The Younger Dryas (YD)
209
4. 4. 5 The Holocene
211
4. 4. 5. 1 The 8.2 k yr event
4. 5 Conclusion
References
212
213
214
Capítulo 5| Long-term and millennial-scale climate variability in north-western France during the last 8
850 years
221
Resumo
222
Résumé
223
Abstract
224
5. 1 Introduction
227
228
230
5. 3. 1 Radiometric dating
230
5. 3. 2 Pollen and dinocyst analyses
231
5. 3. 3 Pollen-based quantitative climate reconstruction
232
5. 4 Results
233
5. 4. 1 Lithostratigraphy and age model
233
5. 4. 2 Evolution of dinocyst assemblages
234
5. 4. 3 Vegetation succession and quantitative climate reconstruction
234
5. 5 Climate variability in north-western France
240
5. 5. 1 Long-term cooling pattern and the Holocene thermal maximum
240
5. 5. 2 Sub-orbital climate variability
243
5. 5. 2. 1 The multi-centennial-scale climate cooling and the 8.2 ka event
244
5. 5. 2. 2 Other possible millennial scale cooling episodes
247
5. 6 Conclusions
247
References
249
Capítulo 6| HOLOCENE CHANGES IN THE DOURO ESTUARY (NORTHWESTERN IBERIA)
257
Resumo
258
Résumé
258
Abstract
259
6. 1 Introduction
261
262
264
265
6. 3. 2 Sedimentological analyses
265
6. 3. 3 Micropalaeontological analyses
266
267
6. 4. 1 Chronology
267
6. 4. 2 Holocene sedimentary processes in the Douro estuary
267
6. 4. 3 Vegetation changes versus variations in pollen catchment area during the Holocene
270
6. 4. 4 Geomorphological changes in the Douro estuary during the Holocene
275
6. 5 Conclusions
277
References
277
Capítulo 7| Conclusions and perspectives
283
Conclusão
293
Índice de Figuras
Capítulo 1|
Fig. I.1 | Variação dos parâmetros astronómicos da terra (excentricidade, obliquidade e precessão dos equinócios)
e da sua resultante, a qual é designada por insolação (W.m -2), durante os últimos 400 000 anos (Berger, 1978). A
insolação representa a quantidade de energia recebida no verão pela terra a 65° N. As listas cinzentas representam
os ciclos interglaciários, os quais são intercalados por períodos glaciários, representados pelas listas brancas.
Fig. I.2 | Parâmetros astronómicos da terra: a) excentricidade, b) obliquidade e c) precessão dos equinócios.
Fig. I.3 | Variabilidade climática orbital nos registos marinhos, continentais e de gelo (adaptado de Tzedakis et al.,
2003). A: curva contínua a negro representa a percentagem de árvores excluindo Juniperus e Pinus e a curva a
tracejado representa a totalidade percentual de árvores; B: curva isotópica do oxigénio contido nas carapaças de
foraminíferos planctónicos e bentónicos da sequência ODP 980 (McManus et al., 1999); C: representa o volume de
gelo versus nível do mar, obtido a partir do δ18O de foraminíferos bentónicos da sequência ODP 980 (McManus et
al., 1999); D: teor em CO2 atmosférico contido na sondagem de gelo de Vostok (Petit et al., 1999); E: curva de
insolação a 40°N e 65°N (Berger, 1978).
Fig. I.4 | Curva de variação da composição isotópica do oxigénio contido na sondagem de gelo GRIP (Dansgaard
et al., 1993). O valor Isotópico do oxigénio (δ18O) contido no gelo representa indirectamente a temperatura
atmosférica do momento no qual ocorreu a acumulação de gelo no pólo Norte (Johnsen et al., 1992). Os números
de 1 a 19 representam os episódios quentes interestadiais. Na parte superior da figura estão representados os
estádios isotópicos marinhos (MIS-Marine isotopic stage) os quais foram definidos a partir da curva de variação do
δ18O contido nas carapaças de foraminíferos bentónicos (Shackleton & Opdyke, 1973). O MIS 3 terminou há cerca
de: a) 24 000 anos segundo Shackleton & Opdyke (1973) e Martinson et al. (1987) ou b) 29 000 anos segundo
Voelker et al. (1998) e van Kreveld et al. (2000). Ao longo do último período glaciário, a variabilidade de D-O é mais
proeminente durante o MIS 3 do que durante o MIS 2 (Voelker et al., 2002) e o MIS 4.
Fig. I.5 | Localização de alguns dos registos paleoclimáticos citados no texto: GRIP (Dansgaard et al., 1993); Vostok
(Blunier et al., 1998; Petit et al., 1999); Byrd (Blunier et al., 1998); V23-81 (Bond et al., 1993; Bond & Lotti, 1995);
ENAM93-21 (Rasmussen et al., 1996; 1997); SU90-24 e SU90-16 (Elliot et al., 1998); SO82-5 (van Kreveld et al., 2000);
NOAMP (Heinrich, 1988) (a sondagem marinha ODP609 localiza-se exactamente na posição da NOAMP, Bond et al.,
1993); HU75-55, 56 e HU87-09 (Andrews & Tedesco, 1992); MD95-2002 (Grousset et al., 2000); PS2644 (Voelker et al.,
1998); ENAM97-09 (Richter et al., 2001); SU90-38 (Cortijo et al., 1995); SU90-03 (Chapman & Shackleton,1998;
Chapman & Maslin,1999; Chapman et al., 2000); A (margem Ibérica): D11957P; SO75-26KL; PO 28-1; PO 8-2; MD952042; SU81-18; MD95-2041; MD95-2040 e MD95-2039) (Lebreiro et al.,1996; Baas et al., 1997; 1998; Zahn et al., 1997;
Abrantes et al., 1998; Bard et al., 2000; Thouveny et al., 2000; de Abreu et al., 2003); RC11-83 (Charles et al., 1996);
TTN057-10/13/21 (Kanfoush et al., 2000). O rectângulo cinza claro representa a localização da cintura de Ruddiman.
Fig. I.6 | Identificação dos eventos de Heinrich na sondagem ODP609 (Bond et al., 1993). Curva de variação: da %
de foraminíferos planctónicos de origem polar (N. pachyderma sin.); de detritos provenientes das calotes glaciárias
do Hemisfério Norte (IRD); de δ18O contido nas carapaças de N. pachyderma sin. e cujos picos representam uma
grande quantidade de fluxo de água fundida (Hemming, 2004).
Fig. I.7 | Calotes glaciárias (Ruddiman, 2001).
Fig. I.8 | Representação esquemática do padrão geral da circulação termohalina (Rahmstorf, 2002). MOC- Atlantic
Meridional Overturning Circulation.
Fig. I.9 | Distribuição global dos registos de D-O durante o MIS 3 (ver Voelker et al., 2002).
Fig. I.10 | Localização das sondagens onde foi efectuado uma correlação directa oceano-continente: 8057B
(Hooghiemstra et al., 1992), ODP 976 (Combourieu Nebout, et al., 1999; 2002), MD95-2042 (Sánchez Goñi et al., 2000;
2002), SO75-6KL (Boessenkool et al., 2001), MD95-2039 (Roucoux et al., 2001; 2005), MD95-2043 (Sánchez Goñi et al.,
2002) e SU81-18 (Turon et al., 2003). Delimitação das zonas biogeográficas (adaptado de Peinado & Rivas-Martinez,
1987).
Fig. I.11 | Alguns registos da variabilidade climática milenar durante o LGIT (adaptado de Goslar et al., 2000): Lago
de Gosciaz na Polónia (Goslar et al., 1992); Bacia de Caríaco (Hughen et al., 1996); GRIP δ18O (Dansgaard et al.,
1993).
Fig. I.12 | Comparação dos registos de GRIP δ18O (Gronelândia) (Dansgaard et al., 1993) com δ18O das sondagens
gelo da Antárctica: Taylor Dome e Byrd e com a variação de Deutério na sondagem de gelo Vostok (adaptado de
Blunier et al., 1998). Esta correlação foi efectuada utilizando a curva de metano de cada um dos registos.
Fig. I.13 | Localização geográfica de alguns dos registos polínicos continentais. a) quadrado A localiza as
sequências de 1 a 5: 1- Laguna de la Roya (Allen et al., 1996), 2- Sanabria March (Allen et al., 1996), 3-Laguna de las
Sanguijuelas (Muñoz Sobrino et al., 2004), 4-Lleguna (Muñoz Sobrino et al., 2004), 5- Pozo do Carballal (Muñoz
Sobrino et al., 1997); b) os pontos 6 a 13 correspondem a: 6- Laguna Lucenza (Santos et al., 2000); 7- Lagoa Lucenza
(Muñoz Sobrino et al., 2001); 8-Lago de Ajo (Allen et al., 1996); 9- Los Tornos (Peñalba, 1994); 10-Saldropo (Peñalba,
1994); 11-Belate (Peñalba, 1994); 12- Atxuri (Peñalba, 1994); 13- Banyoles (Pérez-Obiol & Julià, 1994); c) quadrado B
inclui as sequências 14 a 19: 14- Quintanar de la Sierra (Peñalba et al., 1997); 15- Sierra de Neila - Quintanar de la
Sierra (Ruiz Zapata et al., 2002); 16- Hoyos de Iregua (Gil-Garcia et al., 2002); 17- Laguna Masegosa (Von
Engelbrechten, 1998); 18- Laguna Negra (Von Engelbrechten, 1998); 19- Las Pardillas lake (Sánchez Goñi & Hannon,
1999); 20- Padul (Pons & Reille, 1988); 21- Mougás (Gómez-Orellana et al., 1998); 22- Charco da Candieira (Van der
Knaap & Van Leeuwen, 1995).
Fig. I.14 | Curva de variação do nível global do mar (adaptado de Lambeck et al., 2002).
Fig. I.15 | Localização das zonas biogeográficas (adaptado de Blanco Castro et al., 1997).
Fig. I.16 | Esquema detalhado das principais correntes de superfície do Atlântico Norte: EG-corrente Este da
Gronelândia, Ei-corrente este da Islândia, Gu-Gulf Stream, Ir-corrente de Irminger, La-corrente do Labrador, Nacorrente Norte Atlântica, Nc-corrente do Cap Norte, Ng-corrente da Noroega, Ni-corrente do Norte da Islândia, Pocorrente de Portugal, Sb-corrente de Spitsbergen, Wg-corrente Oeste da Gronelândia. Linhas negras representam as
correntes relativamente quentes enquanto as linhas a tracejado representam correntes relativamente frias
(adaptado de Dietrich et al., 1980).
Fig. I.17 | Representação da localização dos centros de altas e baixas pressões durante o verão e o Inverno ao
longo do Hemisfério Norte (adaptado de Hurrell & Dickson, 2004). As setas representam a direcção dos ventos
dominantes.
Fig. I.18 | Esquema das principais correntes oceânicas que circulam ao longo da margem Ibérica. PCS: Portugal
Current system; ENACWsp: Eastern North Atlantic Central Water de origem sub-polar; ENACWst: Eastern North
Atlantic Central Water de origem sub-tropical; MSW: Mediterranean Sea Water; LSW: Labrador Sea water; NADW:
North Atlantic Deep Water (adaptado de Sprangers et al., 2004).
Fig. I.19 | a) Morfologia da margem continental Ibérica: Canhões submarinos de Mugia (MC), do Porto (PC), de
Aveiro (AC), da Nazaré (NC), de Cascais (CC), de Lisboa (LC), de Setúbal (SC) e de São Vicente (S.VC); montanhas
submarinas de Tore (TS), do Porto (PS), Vasco da Gama (VDGS) e de Vigo (VS).
Fig. I.20 | Morfologia da plataforma continental do noroeste da Península Ibérica (adaptado de Dias et al., 2002).
Fig. I.21 | Morfologia da plataforma continental sudoeste portuguesa (adaptado de Araujo et al., 2002).
Fig. I.22 | Localização do corpo lodoso “Grande Vasière” na plataforma continental Francesa.
Fig. I.23 | Amostras sedimentares de superfície analisadas neste estudo (VIR-18, Ría de Vigo; Laquasup, Estuário do
Douro; PO287-13-2G, Complexo silto-argiloso do Douro; CG11, Complexo silto-argiloso do Minho; MD99-2331, Talude
continental ao largo de Vigo; MD04-2814 CQ, Talude continental ao largo do Porto; Barreiro, Estuário do Tejo; MD992332, Complexo silto-argiloso de Lisboa; FP8-1, Talude continental ao largo de Sines; e MD95-2042, Talude-planície
abissal ao largo de Sines).
Fig. I.24 | Localização das amostras de superfície e das sondagens estudadas ao longo do estuário do Douro. Os
círculos brancos com pinta negra representam as amostras de superfície e os círculos pretos representam as
sondagens.
Fig. I.25 | Localização das sondagens utilizadas neste trabalho. Sondagens marinhas profundas: MD99-2331 e MD032697; sondagem marinha pouco profunda: VK03-58Bis; sondagens estuarinas: Core1 e Core1B.
Capítulo 2|
Fig. II.1 | Fig. 1- Study area. Dashed line divides the Atlantic and Mediterranean biogeographical zones (Blanco
Castro et al., 1997). White circles with a dark point represent the top samples analysed in this study; white circles
represent the modern samples from the European Pollen Database; white circles with a cross represent the studied
cores sites (MD03-2697 and MD99-2331); dark circles represent marine and terrestrial core sites used for comparison
with our study. Continental sequences: a) Square A locates sequences 1 to 5: 1- Laguna de la Roya (Allen et al.,
1996), 2- Sanabria March (Allen et al., 1996), 3-Laguna de las Sanguijuelas (Muñoz Sobrino et al., 2004), 4-Lleguna
(Muñoz Sobrino et al., 2004), 5- Pozo do Carballal (Muñoz Sobrino et al., 1997); b) Sites 6 to 13 correspond to: 6Laguna Lucenza (Santos et al., 2000); 7- Lagoa Lucenza (Muñoz Sobrino et al., 2001); 8-Lago de Ajo (Allen et al.,
1996); 9- Los Tornos (Peñalba, 1994); 10-Saldropo (Peñalba, 1994); 11-Belate (Peñalba, 1994); 12- Atxuri (Peñalba,
1994); 13- Banyoles (Pérez-Obiol and Julià, 1994); c) Square B includes sequences 14 to 19: 14- Quintanar de la Sierra
(Peñalba et al., 1997); 15- Sierra de Neila - Quintanar de la Sierra (Ruiz Zapata et al., 2002); 16- Hoyos de Iregua (GilGarcia et al., 2002); 17- Laguna Masegosa (Von Engelbrechten, 1998); 18- Laguna Negra (Von Engelbrechten, 1998);
19- Las Pardillas lake (Sánchez Goñi and Hannon, 1999); 20- Padul (Pons and Reille, 1988); 21- Mougás (GómezOrellana et al., 1998); 22- Charco da Candieira (Van der Knaap and Van Leeuwen, 1995). The marine cores
represented on the map are: 8057 B (Hooghiemstra et al., 1992), SO75-6KL (Boessenkool et al., 2001), SU81-18 (Turon
et al., 2003) and ODP 976 (Combourieu Nebout, et al., 1999; 2002) and MD95-2039 (Roucoux et al., 2001; 2005).
Fig. II.2 | West to east scheme of the different water masses from the western Iberian margin (adapted from
Sprangers et al., 2004). White circles with a dark point represent southward water flow and white circle with a cross
represent northward water flow. PCS-Portugal Current System; ENACW st-Eastern North Atlantic Central Water of
subtropical origin; ENACW sp-Eastern North Atlantic Central Water of subpolar origin; MSW-Mediterranean Sea Water;
LSW-Labrador Sea Water; NADW-North Atlantic Deep Water.
Fig. II.3 | a) Morphology of the Iberian margin. Location of the surface samples from b) north-western Iberian margin
and c) south-western Iberian margin. White arrows indicate the present-day pattern of pollen dispersion in the
western Iberian margin.
Fig. II.4 | Pollen spectra from western Iberian modern samples. Total temperate and humid (Tot. Temp./Hum.) trees
includes: Alnus, Betula, Corylus, deciduous Quercus and other temperate and humid species (Acer, Fagus, Fraxinus,
Salix, Tilia, Ulmus, Hedera helix, Myrica and Vitis). Total mediterranean (Tot. Mediter.) plants includes: evergreen
Quercus, Olea and Cistus. Taraxacum-type, Asteraceae, Poaceae, Ericaceae and Calluna represent the ubiquist
group. Semi-desert plants include Ephedra, Chenopodiaceae and Artemisia. Climate parameters: Alt: Altitude; PP:
Precipitation; MTCO: Mean temperature of the coldest month; MTWA: Mean temperature of the warmest month;
TANN: Annual temperature.
Fig. II.5 | Pollen assemblages of top samples from coastal and marine western Iberian sites (see also caption of Fig.
II.4). TPC: total pollen concentration.
Fig. II.6 | Galician margin composite record (MD99-2331 and MD03-2697 deep-sea cores). From the left to the right:
corrected radiocarbon ages; marine proxies: δ18O of G. bulloides, % N. pachyderma (s.), ice-rafted detritus (IRD),
Marine and Greenland climatic events; % pollen taxa; pollen zones and chronostratigraphy. Pollen zones were
established using qualitative and quantitative fluctuations of a minimum of 2 curves of ecologically important taxa
(Pons and Reille, 1986). They are defined by the abbreviated name of the core (MD31 or MD97) followed by the
number of the marine isotopic stage (1 or 2) and numbered from the bottom to the top (MD31-2-1 to MD31-2-5 and
MD97-1-1 to MD97-1-6).
Fig. II.7 | - Comparison between continental (Quintanar de la Sierra; Peñalba et al., 1997) and marine (MD99-2331
and MD03-2697) pollen sequences.
Capítulo 3|
Fig. III.1 | Map showing MD99-2331 location and sites of the cores referred in the text: 1: MD95-2040 (Pailler and Bard,
2002; de Abreu et al., 2003; Schönfeld et al., 2003; Narciso et al., 2006), 2: MD95-2039 (Thouveny et al., 2000; Roucoux
et al., 2001; 2005; Schönfeld et al., 2003), 3: PO 28-1 (Abrantes et al., 1998), 4: D11957P (Lebreiro et al., 1996; 1997), 5:
SO75-26KL (Zahn et al., 1997; Boessenkool et al., 2001), 6: PO 8-2 (Abrantes et al., 1998), 7: MD95-2042 (Cayre et al.,
1999; Sánchez Goñi et al., 2000; 2002; Thouveny et al., 2000; Pailler and Bard, 2002), 8: SU81-18 (Bard et al., 2000;
Turon et al., 2003); 9: ODP 976 (Combourieu et al., 2002), 10: SU90-03 (Chapman et al., 2000), 11: ESSCAMP-KS02
(Loncaric et al., 1998; Zaragosi et al., 2001), 12: MD95-2002 (Grousset et al., 2000; Zaragosi et al., 2001), 13: AKS01
(Zaragosi et al., 2001), 14: VM 23-81 (Bond and Lotti, 1995), 15: MD04-2845 (work in progress), 16: SU90-11 (Jullien et
al., in press.), 17: MD03-2705 (Jullien et al., submitted), 18: OCE326-GGC5 (McManus et al., 2004).
Fig. III.2 | Chronostratigraphy of the MD99-2331 record. AMS radiocarbon dates are represented by triangles while
the calibrated ones are represented by squares. North Atlantic Heinrich events are delimited by both the age limits
(not calibrated) proposed by Elliot et al. (2002) and by the results obtained from the multi-proxy study of the MD992331 record (see below). White triangles and squares reflect the rejected levels for the model age while the dark
ones represent the accepted ages.
Fig. III.3 | Comparison between long term trends of MD99-2331 record and Greenland temperatures (Sánchez Goñi
et al., in prep.) during the Late MIS 3 and MIS 2 against age (cal yr BP). From the bottom to the top: MD99-2331
benthic δ13C; percentages of Pinus, temperate and humid trees and Poaceae and Greenland temperatures. Dashed
line represents the long term trend of each signature.
Fig. III.4 | Multi-proxy results of MD99-2331 record. From bottom to top: ice-rafted detritus (IRD) concentrations,
percentages of planktonic foraminifera associations (polar, sub-polar and warm), planktonic foraminifera-based
winter and summer SST estimates, alkenone-based annual SST reconstruction, δ18O of G. bulloides, percentages of
temperate and humid trees, Pinus percentages and Greenland temperatures (Sánchez Goñi et al., in prep.). Grey
lines represent the Heinrich events. Note: the age limits of H4 that are based on GISP2 chronology are slightly different
from those estimated from the calibration using NGRIP.
Fig. III.5 | Response of the north-western Iberia vegetation to the complex pattern of Heinrich events in the Iberian
margin. From bottom to top: Heinrich events, percentages of Pinus, percentages of Calluna representing wet
conditions, percentages of semi-desert plants reflecting continental dryness and total of pollen concentration.
Dashed line separated wet from dryness conditions during H events.
Fig. III.6 | Prevailing NAO negative conditions scheme (adapted from Wanner et al., 2001).
Fig. III.7 | Prevailing NAO positive conditions scheme (adapted from Wanner et al., 2001).
Capítulo 4|
Fig. IV.1 | Study area. Location of deep-sea cores referred in the text: OCE326-GGC5 (McManus et al., 2004); MD952002 (Zaragosi et al., 2001; Auffret et al., 2002; Ménot et al., 2006) and MD99-2331 (Naughton et al., 2006; in prep.).
Fig. IV.2 | Pollen diagram of MD03-2697 deep-sea core against depth. From left to right: calibrated ages and
percentages of selected pollen taxa. The stratigraphy is based on previous work by (Naughton et al., 2006) where the
Oldest Dryas represents the continental counterpart of Heinrich 1 event in the ocean.
Fig. IV.3 | Multi-proxy record of MD03-2697 against calibrated ages. From bottom to top: percentages of selected
pollen taxa (trees: Betula, Corylus, deciduous Quercus, Pinus;
Calluna and semi-desert plants: Artemisia,
Chenopodiaceae and Ephedra); Ice-rafted detritus (IRD); planktonic foraminifera associations; Sea Surface
Temperature (SST) estimates; δ18O of planktonic and benthic foraminifera and Greenland temperatures (Sánchez
Goñi et al., in prep.).
Fig. IV.4 | Long-term and small-scalle pattern of vegetation changes in north-western Iberia. From bottom to top:
benthic foraminifera δ18O; percentages of temperate trees includes (Acer, Alnus, Betula, Corylus, Cupressaceae,
deciduous Quercus, Fagus, Fraxinus excelsior-type, Pinus, Quercus ilex-type, Salix, Tilia and Ulmus); percentages of
Pinus, percentages of deciduous Quercus and summer insolation at 45° N (after Berger, 1978).
Capítulo 5|
Fig. V.1 | Location of shelf core VK03-58Bis; and deep-sea core MD99-2551 (Ellison et al., 2006).
Fig. V.2 | Lithology and synthetic pollen diagram against depth (cm). From left to right: radiocarbon and calibrated
ages; lithology (after Folliot, 2004) including T. communis level (represented by small shells); dinocyst percentages
(Operculodnium centrocarpum; Total of Spiniferites and Lingulodinium machaerophorum); pollen diagram and
pollen zones.
Fig. V.3 | Pollen diagram and quantitative pollen-based climate estimates against depth (cm). From left to right:
calibrated ages; selected pollen taxa from the synthetic pollen diagram (other deciduous trees include: Fraxinus
excelsior-type, Tilia and Ulmus); climate parameters: PANN (mean annual precipitation); difference between the
temperature of the warmest (MTWA) and the coldest (MTCO) months (seasonality) and TANN (mean annual
temperatures). Dashed lines represent maxima (bold) and minima values and the dark line represents mean values.
Grey dashed lines represent the tendency of each curve; pollen zones.
Fig. V.4 | Correlation between vegetation changes, quantitative climate estimates, summer insolation at 45° N and
precessional signal (after Berger, 1978) and δ18O-isotope composition of the NorthGRIP ice-core (Johnsen et al.,
2001) during the Holocene. Temperate and humid trees include: Acer, Alnus, Betula, Corylus, Cupressaceae,
deciduous Quercus, Fagus, Fraxinus excelsior-type, Pinus, Quercus ilex-type, Salix, Tilia and Ulmus while
Brassicaceae, Caryophyllaceae, Asteraceae (including Aster- and Anthemis- types) and Taraxacum-type,
Cyperaceae, Ericaceae and Calluna, Plantago, Poaceae and semi-desert plants (including Chenopodiaceae,
Artemisia and Ephedra) are integrated in the herbaceous plants association.
Fig. V.5 | Correlation between selected pollen taxa, quantitative climate estimates (PANN, TANN, MTCO, MTWA and
seasonality) and δ18O-isotope composition of the NorthGRIP ice-core (Johnsen et al., 2001) during the Holocene. The
8.2 kyr event is represented by the dark grey bar which is superimposed to 8.6-8.0 kyr event represented by the light
grey bar. Dark arrows indicate possible millennial-scale cooling events during the Holocene.
Fig. V.6 | Present day and past marine biogeographical zones in the North-East Atlantic (adapted from Funder et al.,
2002). Bold dashed lines represent the limits of the present-day marine biogeographical zones in the North-East
Atlantic; Grey dashed lines represent: a) the northward displacement of the boreal southern limit during the early
Eemian (Funder et al., 2002) and b) the southward displacement of the boreal southern limit during the during 8.6-8.0
kyr event (this work).
Capítulo 6|
Fig. VI.1 | a) Douro estuary localisation in the Iberian Península. b) Core (1, 1B and 2) and surface sampling sites.
Dark circles represent core sites and white circles surface samples sites. Dark lines represent the palaeoisobathic
curves defined by Carvalho and Rosa (1988). Dashed line represents the ancient direction of the river main channel
flow and bold dark line the present day river main channel flow. This palaeobathymetric map shows the
palaeovalley of the Douro river.
Fig. VI.2 | Lithlogy of the Douro sequence. Depth is represented in Ordnance Datum which is a coordinate system for
heights above mean sea level (optometric heights). Five calibrated ages are also represented along the two cores.
Fig. VI.3 | Lithology of the Douro sequence. Depth is represented in Ordnance Datum which is a coordinate system
for heights above mean sea level (optometric heights). Five calibrated ages are also represented along the two
cores.
Fig. VI.4 | a) Grain roundness and b) effective sphericity of gravel pebbles plotted against particle size. Two black
lines delimit the different roundness and sphericity averages defined by Dobbkins and Folk (1970), representing the
average limit of grain characterising river and/or low and high-energy beach environments. Dark squares represent
all the measures obtained and Circle represents the value means of all measures.
Fig. VI.5| Pollen diagram. From the left to the right: lithology, calibrated ages, arboreal pollen, AP (total of arboreal
pollen), Pinus, NAP (total of non-arboreal pollen), pollen of herbaceous plants and pollen zones.
Fig. VI.6| LAQUASUP: surface samples sites and pollen percentages.
Fig. VI.7| Conceptual model of the Holocene geomorphological evolution of the Douro estuary: a) between 10720
and 6530 cal yr BP, b) from 6530 to 1500 cal yr BP and c) the last 1500 years.
Índice de Tabelas
Capítulo 1 |
Tab. I.1 | Localização das amostras de superfície. Da esquerda para a direita encontra-se representado o nome das
amostras, a profundidade na coluna sedimentar, a latitude, a longitude, a profundidade da coluna de água, o ano
da colheita das amostras, o nome dos projectos científicos ou o nome das missões oceanográficas ou o nome das
Instituições que forneceram as amostras.
Capítulo 2|
Tab. II.1| Radiocarbon ages of MD99-2331 and MD03-2697 deep-sea cores. a Not acceptable dating (bioturbated
layers); b Radiocarbon dates too old (not used); c dates calibrated by matching conventional AMS 14C with calendar
ages estimated for MD95-2042 deep-sea core by Bard et al. (2004).
Tab. II.2| Location, water depth and year of sample sampling from coastal, shelf and slope sequences of the Iberian
margin.
Tab. II.3| Description of the pollen zones in the Galician margin composite core and respective chronostratigraphy.
Tab. II.4| Description of pollen zones from the well-dated reference sites of Quintanar de la Sierra (Peñalba 1994,
Peñalba et al., 1997), Laguna de la Roya (Allen et al., 1996) and Padul (Pons and Reille, 1988).
Tab. II.5| Holocene tree succession in north-western Iberia.
Capítulo 3|
Tab. III.1| Radiocarbon ages of MD99-2331 deep sea cores. Bold levels represent the accepted ages while not bold
ones represent the rejected ages for the age model.
Capítulo 4|
Tab. IV.1| Radiocarbon ages of MD03-2697 deep-sea core and one level from the twin core (MD99-2331).
Radiocarbon dates too young or too old and b Not acceptable dating (bioturbated layers).
a
Capítulo 5|
Tab. V.1| Radiocarbon ages from VK03-58Bis and VK03-58 and VK03-59Bis shelf cores.
Capítulo 6|
Tab. VI.1 | Radiocarbon and calibrated dates from the site under study.
Índice de Anexos
Anexo A
Climate variability of the last five isotopic interglacials: direct land-sea-ice correlation from the
multiproxy analysis of north western Iberian margin deep-sea cores.
S. Desprat, M.F., Sánchez Goñi, F., Naughton, J.-L., Turon, J. Duprat, B. Malaizé, E. Cortijo and J.-P.
Peypouquet. in press in The climate of past interglacials. Elsevier publications.
Paleoenvironmental evolution of estuarine systems during the last 14000 years – the case of Douro
estuary (NW Portugal).
T. Drago, C.Freitas, F.Rocha, M.Cachão, J.Moreno, F.Naughton, C.Fradique, F.Araújo, T.Silveira, A.Oliveira,
J.Cascalho, F.Fatela. (in press). In press in Journal of Coastal Research, SI 39, 7 pp.
Capítulo 1 | Introdução
Face à crescente inquietude relativa ao aquecimento actual do
planeta, torna-se necessário compreender a dinâmica natural do clima no
passado, assim como identificar as variações ambientais que dela
resultaram.
As variações climáticas são um fenómeno global as quais resultam da
interacção entre a atmosfera, hidrosfera, criosfera, litosfera e biosfera. Estes
diferentes componentes interagem por sua vez com o clima produzindo
retroacções negativas ou positivas ou, pelo contrário, amplificam ou
reduzem um dado sinal climático.
Os
mecanismos
forçadores
externos
(os
quais
se
encontram
intimamente ligados à posição da terra em relação sol e à constante solar) e
internos (circulação termohalina, mecanismo interno ao gelo, erupções
vulcânicas, concentração de CO2, albedo, etc.....) são responsáveis pela
frequência, duração e amplitude das variações climáticas.
De forma a detectar e compreender a frequência, duração e
amplitude, assim como os mecanismos responsáveis pela variabilidade
climática natural, torna-se necessário reconstituir, para o passado, o impacto
desta variabilidade nos cinco sub-sistemas climáticos, no passado, utilizando
uma cronologia única. Esta reconstituição pode ser realizada através da
correlação directa entre registos paleoclimáticos marinhos, continentais e de
gelo.
A correlação indirecta entre sequências sedimentares terrestres,
marinhas e de gelo é no entanto difícil devido à falta de precisão existente
entre a conexão dos diferentes modelos de idade. A maioria dos registos
climáticos
encontram-se
dispersos
geograficamente
pelo
mundo
e
apresentam modelos de idades distintos os quais se baseiam em diversos
tipos de metodologias tais como: datações radiométricas (14C, U/Th, etc),
contagem das camadas de gelo anuais nas sondagens da Gronelândia, e
ainda com base em modelos físicos de acumulação de gelo na Antárctica.
De forma a contornar esta problemática, o estudo de sondagens
marinhas ricas em pólen e esporos permite-nos obter informações sobre a
história da vegetação, assim como do clima no continente adjacente. Este
tipo
de
sondagens
englobam
ainda
uma
série
de
indicadores
paleoclimáticos marinhos que permitem estimar a temperatura e salinidade
da
massa
oceânica
de
superfície
(foraminíferos
planctónicos,
dinoflagelados, nanoplancton, diatomáceas, etc), as condições da massa
oceânica de fundo (Mg/Ca, ostracodos, foraminíferos bentónicos, δ13C), a
dinâmica dos icebergues e a instabilidade das calotes polares (sedimentos
grosseiros) assim como o volume de gelo acumulado nos pólos (δ18O de
foraminíferos bentónicos). A comparação directa entre os diferentes tipos de
registos paleoclimáticos permite-nos identificar as variações climáticas que
afectaram o continente e correlacioná-las directamente com a resposta dos
outros componentes do sistema climático (oceano, gelo e atmosfera). É-nos
possível ainda documentar eventuais não contemporaneidades entre a
resposta dos vários sub-sistemas climáticos a uma dada variação climática,
assim como compreender a frequência e a natureza dessas variações. A
comparação entre dados e resultados obtidos a partir de simulações
numéricas permite-nos ainda testar a validade das últimas assim como
confirmar os mecanismos utilizados nessas mesmas estimativas.
Nos últimos milhões de anos, o clima foi sujeito a alternâncias entre
períodos glaciários e interglaciários, segundo ciclos de cerca de 100 000
anos, induzidos por variações na insolação de verão. Super imposta a esta
ciclicidade orbital, outras oscilações da ordem do milénio foram detectadas
em sondagens de gelo da Gronelândia. Estas oscilações, ditas DansgaardOeschger (D-O), são marcadas por uma alternância entre aquecimentos
súbitos e arrefecimentos progressivos. Alguns dos episódios frios são
considerados eventos extremos e resultam da forte descarga de icebergues
no Atlântico Norte.
Estes eventos, ditos Heinrich, foram identificados em várias sequências
marinhas do Atlântico Norte, nomeadamente na cintura de Ruddiman entre
40° e 55° N, e são caracterizados pela forte presença de grãos detríticos de
dimensões superiores a 150 μm (designados por IRD - Ice Rafted Detritus),
assim como por um sinal magnético bastante importante. Estas camadas de
IRDs foram ainda detectadas fora da sua zona preferencial de acumulação,
2
F. Naughton, 2007
nomeadamente nas médias latitudes do Atlântico Norte, incluindo na
margem Ibérica.
Nos
últimos
40
000
anos,
nomeadamente H4 (34.9-33.9 k anos
(22.1-20.4 k anos
14C
quatro
14C
eventos
de
Heinrich
BP), H3 (27.4-26.1 k anos
BP) e H1 (15.1-13.4 k
14C
14C
(H),
BP), H2
anos BP) foram registados no
Atlântico Norte. No entanto, estes eventos apresentam um padrão
complexo, composto por duas fases distintas, no sudoeste da margem
ibérica (Bard et al., 2000; Sánchez Goñi et al., 2000). A primeira fase é
caracterizada por uma diminuição abrupta da temperatura das águas de
superfície (SST-Sea Superfície Temperatura), uma fraca presença de IRDs e
um clima relativamente húmido no continente adjacente, enquanto que a
segunda fase é caracterizada pela deposição máxima de IRDs e por um
clima árido no continente. Este padrão complexo não foi até à data
documentado para o noroeste da Península Ibérica. Várias hipóteses foram
propostas na tentativa de explicar este padrão complexo associado aos
eventos de Heinrich (Bard et al., 2000; Sánchez Goñi et al., 2000; Abrantes et
al., 1998). No entanto, várias questões permanecem ainda sem resposta.
Por essas razões um dos principais objectivos deste trabalho seria
descrever com precisão o impacto destes eventos extremos no noroeste da
margem Ibérica e continente adjacente, assim como discutir sobre os
eventuais mecanismos responsáveis pelo padrão complexo deixado pelos
eventos de Heinrich nesta região.
Neste trabalho, pretendemos ainda compreender as interacções
existentes entre os diferentes sub-sistemas climáticos durante o último
máximo glaciar (LGM - Last Glacial Maximum). Sabe-se que a vegetação
respondeu contemporaneamente às variações de SST que caracterizam as
oscilações de D-O durante a fase tardia do estádio isotópico marinho 3 (late
MIS3) (Sánchez Goñi et al., 2000; 2002). A expansão da floresta temperada é
geralmente associada a um aquecimento oceânico superficial, enquanto
que a regressão desta floresta corresponde a um arrefecimento das
condições oceânicas de superfície. Contudo, durante o período de máxima
extensão das calotes glaciárias no Hemisfério Norte, observa-se uma
assincronia entre a SST (elevada nas médias latitudes do Atlântico Norte
oriental) (e.g Chapman et al., 2000; de Abreu et al., 2003; Morey et al., 2005),
3
F. Naughton, 2007
e a biosfera continental adjacente (dominada por uma paisagem aberta
indicando temperaturas frias no continente) (Peyron et al., 1998). Por essas
razões,
o
segundo
objectivo
deste
trabalho
foi
então,
de
tentar
compreender esta situação paradoxal.
Pretendemos ainda identificar e compreender a resposta da
vegetação das médias latitudes do Atlântico Norte aos eventos sucessivos
que caracterizam a última deglaciação, nomeadamente: o final do LGM, o
H1, o evento quente Bölling-Alleröd (interestadial 1 nas sondagens de gelo
da Gronelândia) e o evento frio Dryas recente (estadial 1 nas sondagens de
gelo da Gronelândia), assim como o aquecimento que caracteriza o início
do Holocénico. A comparação detalhada entre sequências polínicas
marinhas e continentais permitiu-nos correlacionar claramente a clássica
estratigrafia continental com os eventos oceânicos detectados no Atlântico
Norte e os eventos registados nas sondagens de gelo da Gronelândia.
Finalmente, interessamo-nos em identificar e compreender a resposta
da vegetação face à variabilidade climática orbital e sub-orbital que
caracteriza o actual interglaciário (Holocénico).
Pretendeu-se ainda detectar e definir o intervalo de tempo associado
ao Máximo Térmico Holocénico (HTM) na Europa centroeste e sudoeste.
De forma a alcançar os objectivos mencionados foram utilizadas três
sondagens às quais foi aplicado um estudo multidisciplinar (indicadores
paleoclimáticos continentais, marinhos e de volume de gelo acumulado nos
pólos) de alta resolução temporal (~100 anos).
Previamente a efectuar a correlação directa oceano-continente-gelo
dessas sondagens, foi efectuada uma calibração do sinal polínico marinho
ao longo da margem Ibérica. De facto, a comparação do sinal polínico
actual de amostras colhidas na margem Ibérica (num perfil de sul para
Norte) com amostras continentais (representantes das duas regiões
biogeográficas
provenientes
principais:
da
base
Mediterrânica
de
dados
a
sul
e
Europeia
Atlântica
a
norte)
(http:/www.imep-
cnrs.com/pages/EPD.htm; Peyron et al., 1998; Barboni et al., 2004) permitenos verificar:
- se assinatura polínica marinha actual representa uma imagem
integral da vegetação regional do continente adjacente, e;
4
F. Naughton, 2007
- se a assinatura polínica das amostras de superfície costeiras e
marinhas do noroeste e sudoeste da Península Ibérica é semelhante aos
espectros polínicos actuais representantes das regiões biogeográficas
Atlântica e Mediterrânica, respectivamente.
De forma a compreender melhor o sinal polínico ao longo desta
margem foram ainda determinados os padrões de transporte e dispersão
polínica, do continente para o mar aberto.
As sondagens marinhas utilizadas ao longo deste trabalho, a MD992331 (42° 09’ 00 N, 09° 40’ 90 W; 2110 m de profundidade) e a MD03-2697 (42°
09’ 59 N, 59° 42’ 10 W; 2164 m de profundidade), foram recolhidas no
noroeste da Península Ibérica pelo navio oceanográfico Marion Dufresne
durante as campanhas GINNA e GEOSCIENCES. No entanto, a resolução
temporal destas sondagens para o Holocénico é bastante fraca pelo que,
foi utilizada uma terceira sondagem, a Vk03-58Bis, recolhida na plataforma
continental noroeste francesa, na “Grande Vasière” (47°36’ N, 4°08’ W; 98 m
de profundidade), a qual é constituída por 2.72 m de sedimento. Os
principais indicadores climáticos utilizados neste trabalho foram: os grãos de
pólen, os sedimentos grosseiros provenientes da fusão dos icebergues, os
foraminíferos planctónicos e os isótopos de oxigénio de foraminíferos
bentónicos.
O último objectivo deste trabalho seria documentar e compreender o
impacto da variabilidade climática Holocénica na evolução dos sistemas
costeiros do norte de Portugal. Sabe-se que o clima tem uma forte influência
na evolução dos sistemas costeiros. De facto, a variabilidade climática
orbital induz oscilações no volume de gelo acumulado nos pólos que, por
sua vez provocam modificações do nível do mar as quais vão agir na
evolução dos sistemas costeiros. A relação entre as variações do nível do
mar e a evolução dos sistemas costeiros foi aprofundada para a zona sul
portuguesa (Freitas et al., 2002 ; 2003).
De forma a compreender o impacto das variações climáticas e
consequentes modificações do nível do mar na evolução geomorfológica
do estuário do Douro (Norte de Portugal) durante o Holocénico, foi
efectuado um estudo polínico e sedimentar em duas sondagens estuarinas
5
F. Naughton, 2007
colhidas no estuário do Douro. Neste trabalho foi ainda efectuada a
distinção dos diferentes mecanismos responsáveis pela evolução deste
sistema costeiro.
INTRODUCTION
Face à l’inquiétude croissante sur le réchauffement actuel de la
planète dû sans doute au moins en partie aux activités humaines, il est
nécessaire plus que jamais de connaître la dynamique naturelle du climat
dans le passé et les modifications environnementales qui en découlent. La
compréhension du système climatique naturel peut permettre de dissocier ce
qui dans le changement climatique actuel est imputable à l’homme de ce
qui est propre à la cyclicité naturelle.
Le changement climatique est un phénomène global qui agît sur
l’atmosphère, l’hydrosphère, notamment les océans, la cryosphère, en
particulier les calottes de glace polaires, la lithosphère et la biosphère. Ces
différents réservoirs agissent, à leur tour, sur le climat en produisant des
rétroactions négatives ou, au contraire, amplifiant le signal climatique. Le
forçage externe lié à la position de la Terre par rapport au Soleil et à la
constante
solaire
ainsi
que
plusieurs
forçages
internes
(circulation
thermohaline, mécanique interne de la glace, éruptions volcaniques,
concentration du CO2, albédo…) sont les responsables de la fréquence, la
durée et l’amplitude du changement climatique. Pour comprendre ces trois
paramètres et les mécanismes associés au changement climatique naturel il
est donc nécessaire de reconstituer son impact sur ces cinq réservoirs dans le
passé avec une chronologie commune.
Cette reconstitution peut être réalisée par la corrélation des
enregistrements paléoclimatiques marins, de glace et continentaux. Toutefois
la corrélation des séquences terrestres avec les archives marines et de glace
est difficile en raison du manque de précision des différents calages
chronologiques d’une séquence à l’autre. En effet, ces séquences dispersées
géographiquement, ont de plus des modèles d’âge différents, certains basés
sur des datations radiométriques (14C, U/Th…), d’autres sur le comptage des
6
F. Naughton, 2007
couches de glace accumulées annuellement sur le Groenland et encore
d’autres
sur
des
modèles
physiques
d’accumulation
de
glace
en
Antarctique. Une façon de contourner ce problème est de travailler sur des
carottes marines à sédimentation continue et non perturbée riches en pollen
et spores qui nous renseignent sur l’histoire de la végétation et, en
conséquence, du climat du proche continent. Ces carottes ont de plus
l’avantage de renfermer des indicateurs paléoclimatiques provenant
d’autres réservoirs permettant d’estimer les températures et la salinité des
eaux de surface (foraminifères planctoniques, kystes de dinoflagellés,
alcénones, nannoplancton, diatomées), les conditions des eaux du fond
(Mg/Ca, ostracodes, foraminifères benthiques, δ13C), la dynamique des
icebergs et l’instabilité des calottes polaires (sédiments grossiers) et le volume
de glace stocké aux pôles (isotopes des foraminifères benthiques). La
corrélation
directe
de
ces
différents
types
d’enregistrements
paléoclimatiques permet d’une part d’identifier les changements climatiques
qui ont affecté le continent et de les corréler directement avec la réponse
d’autres composants du système climatique (océan, glace, atmosphère). On
pourra ensuite documenter d’éventuels déphasages dans la réponse de ces
réservoirs à un même changement climatique et enfin discuter de la
fréquence et nature de ces changements. Une comparaison entre données
et sorties des modèles permet en plus de tester la validité de ces dernières et,
donc, la robustesse des mécanismes impliqués dans les différents modèles
pour reproduire le changement climatique. En particulier, il est possible de
tester le rôle possible des changements de la végétation sur le climat.
Au cours du dernier million d’années, le climat a été rythmé par des
alternances entre périodes glaciaires et interglaciaires, sur des cycles de
100 000 ans environ, induites par les variations d’insolation. Surimposée à
cette cyclicité orbitale de long terme, d’autres oscillations d’ordre millénaire
ont été détectées, au cours de la dernière période glaciaire, dans des
carottes de glace du Groenland. Ces
Oeschger
(D-O),
sont
marquées
oscillations, dites de Dansgaard-
par
une
alternance
entre
des
réchauffements importants très rapides suivis de refroidissements progressifs.
Certains de ces événements sont associés à des refroidissements extrêmes
des eaux de surface de l’Atlantique Nord induites par l’introduction de
7
F. Naughton, 2007
grandes quantités d’eau de fonte provenant des importantes débâcles
d’icebergs dans l’hémisphère Nord. Ces événements, dits d’Heinrich, sont
identifiés dans les séquences marines de l’Atlantique Nord, et notamment
dans la ceinture de Ruddiman (45°-55° N). Ils sont caractérisés par une forte
accumulation des grains détritiques de dimension supérieure à 150 μm
(désignés IRD – Ice Rafted Detritus), ainsi que par un signal magnétique
robuste. Des couches d’IRD ont été détectées aussi en dehors de cette zone
préférentielle de déposition, et en particulier dans des séquences marines
des moyennes latitudes incluant celles de la marge ibérique.
Pour la période qui nous intéresse, les derniers 40 000 ans, quatre
événements d’Heinrich H4 (34.9-33.9
H2 (22.1-20.4
14C
14C
k ans BP), H3 (27.4-26.1
k ans BP) et H1 (15.1-13.4
14C
14C
k ans BP),
k ans BP) on été enregistrés
dans cette région.
Les événements d’Heinrich montrent des scénarii complexes, avec
deux phases distinctes enregistrées dans les carottes marines de la marge
sud-ouest de l’Ibérie (Bard et al., 2000 ; Sánchez Goñi et al., 2000). La
première phase correspond à une forte diminution de la température des
eaux de surface (SST-Sea Surface Temperature), la presque absence d’IRD et
un climat relativement humide sur le continent. La deuxième phase est
caractérisée par le dépôt maximal d’IRD, un climat aride sur le continent
adjacent tout en conservant des SST froides. Les changements associées à
ces phases dans le nord-ouest de la Péninsule Ibérique n’ont pas été,
toutefois, documentés. Quelques hypothèses ont été proposées pour
expliquer ces scénarii complexes associés aux événements d’Heinrich.
Néanmoins des questions restent en suspend.
C’est pour cela qu’un des principaux objectifs de ce travail a été,
dans un
premier temps, de décrire avec précision l’impact que ces
événements ont produit sur le nord-ouest de la marge ibérique et le continent
adjacent. Ensuite, nous avons discuté des mécanismes associés à ces
scénarii.
D’autre part, ce travail s’est aussi intéressé à la compréhension des
interactions entre les différents réservoirs terrestres pendant le dernier
maximum glaciaire (LGM – Last Glacial Maximum). Généralement, la
végétation a répondu de façon synchrone aux variations D-O des SST
8
F. Naughton, 2007
(Sánchez Goñi et al., 2000 ; 2002). L’expansion de la forêt caducifoliée est
associée à des réchauffements océaniques importants, inversement le retrait
de cette forêt correspond à un refroidissement de la surface marine.
Cependant, durant la période d’extension maximale des calottes glaciaires
dans l’hémisphère nord, nous avons observé un découplage entre les SST, qui
sont élevées dans les moyennes latitudes de l’Atlantique Nord oriental (e.g
Chapman et al., 2000 ; de Abreu et al., 2003 ; Morey et al., 2005), et la
biosphère du continent adjacent où les paysages ouverts indiquent des
températures froides (Peyron et al., 1998). Comprendre cette situation
paradoxale a été le deuxième objectif de cette thèse.
De plus, nous nous sommes aussi attachés à identifier et comprendre la
réponse de la végétation des moyennes latitudes aux événements successifs
qui caractérisent la dernière déglaciation : la fin du LGM, l’H1, l’événement
chaud du Bölling-Alleröd (interstade 1 des carottes Groenlandaises), et
l’événement froid du Dryas Récent (stade 1 des carottes Groenlandaises)
ainsi que le réchauffement qui caractérise le début de l’Holocène. Une
comparaison détaillée entre séquences polliniques marines et continentales
nous a permis de corréler sans ambigüité la stratigraphie continentale
classique avec les événements océaniques et groenlandais.
Enfin, nous avons aussi considéré la variabilité climatique orbitale et
sub-orbitale qui caractérise l’actuel interglaciaire (l’Holocène). Comment la
végétation
des
moyennes
latitudes
a
réagi
t’elle
aux
forçages
astronomiques ?, Quelle est sa réponse aux événements froids de l’Atlantique
Nord (l’événement multi-centenaire 8.6-8.0 k ans et l’événement de 8.2 k
ans), associés aux phases finales de purge des lacs d’Agassiz et d’Ojibway ?
Pour atteindre les objectifs mentionnés ci-dessus, nécessitant d’une
chronologie commune et fiable entre les différents réservoirs terrestres, nous
avons utilisés des carottes marines riches en pollen. Forts de cette approche
de corrélation directe entre indicateurs climatiques marins, terrestres et de
glace, nous avons appliqué une analyse à très haute résolution (~100 ans)
afin de détecter le plus grand nombre d’événements climatiques des
derniers 40 000 ans.
Avant d’effectuer la corrélation directe océan-continent-glace, nous
avons tout d’abord calibré le signal pollinique marin de la zone d’étude.
9
F. Naughton, 2007
Nous avons confronté le signal pollinique actuel des échantillons de surface
prélevée sur la marge (du Sud vers le Nord) avec ceux de la Péninsule
Ibérique (représentant les deux principales régions biogéographiques
Méditerranéenne au sud et Atlantique au nord) provenant de la base de
données polliniques européenne (Peyron et al., 1998; Barboni et al., 2004;
http:/www.imep-cnrs.com/pages/EPD.htm ), de façon à montrer que :
- les assemblages polliniques de la marge Ibérique représentent une
image intégrée de la végétation régionale qui colonise le continent
adjacent ;
-
les
communautés
biogéographiques
forestières
pré-citées
sont
bien
présentes
sur
les
discriminées
par
deux
les
zones
spectres
polliniques marins du sud et du nord de la marge Ibérique, respectivement.
Pour améliorer la compréhension du signal pollinique sur cette marge,
les scénarii de dispersion pollinique actuelle et les principaux mécanismes
responsables du transport sur cette région ont été approfondis.
Les carottes marines utilisées, MD99-2331 (42° 09’ 00 N, 09° 40’ 90 W; 2110 m
profondeur) et MD03-2697 (42° 09’ 59 N, 59° 42’ 10 W; 2164 m profondeur), ont
été prélevées par le bateau océanographique Marion Dufresne en face de
la Galice (nord-ouest de la Péninsule Ibérique) lors des campagnes GINNA et
GEOSCIENCES. Toutefois, la résolution faible de ces carottes pour étudier la
variabilité climatique des derniers 10 000 ans (Holocène) nous a amené à
travailler sur une autre carotte marine, cette fois-ci moins profonde, prélevée
en face de Brest dans la Grande Vasière (nord-ouest de la France, 47°36’ N
et 4°08’ W) et constituée de 2.72 m de sédiment.
Les indicateurs climatiques (proxies) utilisés ont été principalement : le
pollen, la fraction de sédiment grossier provenant de la fonte des icebergs,
les foraminifères planctoniques et les isotopes de l’oxygène des foraminifères
benthiques.
Ces
proxies
vont
documenter
les
variations
dans
les
communautés végétales et des foraminifères indiquant la variabilité
climatique sur le continent et l’océan, respectivement ; les périodes des
débâcles et fonte d’icebergs et la variation du volume de la calotte polaire
de l’Hémisphère Nord.
10
F. Naughton, 2007
Un but supplémentaire de notre thèse était de déceler l’impact des
variations climatiques sur l’évolution des systèmes côtiers du nord-ouest de la
marge Ibérique au cours des derniers 10 000 ans. Nous savons que le climat a
une très forte influence sur l’évolution de ces systèmes. En particulier, la
variabilité orbitale du climat est traduite par des variations du volume de
glace accumulé aux pôles qui induisent des modifications du niveau marin
avec un impact important sur l’évolution des systèmes côtiers. La relation
entre les variations du niveau marin et l’évolution des systèmes côtiers de la
marge occidentale ibérique a été étudiée pour le sud de la Péninsule
ibérique (Freitas et al., 2002 ; 2003). Cependant, la réponse des systèmes
côtiers du nord-ouest ibérique reste mal connue. Afin de comprendre
l’évolution géomorphologique de cet environnement pendant l’Holocène
nous avons réalisé les analyses pollinique et sédimentaire d’une carotte
estuarienne prélevée dans l’estuaire du Douro. Dans ce travail, les
mécanismes de forçage global lié aux variations du niveau marin sont
distingués des mécanismes locaux qui ont agît sur l’évolution de ce système
côtier.
11
F. Naughton, 2007
1. 1 Paleoclimatologia do quaternário recente, contexto e objectivos
principais deste trabalho
1. 1. 1 Variabilidade climática orbital
Nos últimos milhões de anos o clima terrestre sofreu variações de
longo-termo, entre condições glaciárias e interglaciárias, como resposta a
modificações físicas que ocorreram entre o sol e a terra (Imbrie et al., 1992).
Este mecanismo forçador externo, resulta da associação dos parâmetros
astronómicos (excentricidade, obliquidade e precessão dos equinócios), que
controlam a distribuição sazonal e latitudinal da energia proveniente do sol
(Insolação) (Fig. I.1) (Berger & Loutre, 2004).
Fig. I.1 | Variação dos parâmetros astronómicos da terra (excentricidade, obliquidade e precessão dos
equinócios) e da sua resultante, a qual é designada por insolação (W.m -2), durante os últimos 400 000
anos (Berger, 1978). A insolação representa a quantidade de energia recebida no verão pela terra a 65°
N. As listas cinzentas representam os ciclos interglaciários, os quais são intercalados por períodos
glaciários, representados pelas listas brancas.
A órbita da terra à volta do sol varia de circular a elíptica, segundo
ciclos próximos de 100 000 e 400 000 anos (Fig. I.2.a). O grau de
achatamento da elipse em relação ao círculo caracteriza a excentricidade.
O valor da excentricidade varia entre 0 a 0.05 (Berger & Loutre, 2004). Este
pequeno valor da excentricidade da órbita terrestre afecta fracamente a
quantidade de energia solar recebida pela terra anualmente.
A obliquidade representa a relação entre o ângulo de inclinação do
eixo da terra e a perpendicular ao plano da sua órbita (plano da eclíptica),
a qual vai afectar a quantidade de energia recebida pela terra durante as
estações do ano (Fig. I.2.b). O valor da obliquidade varia entre 22° e 25°
segundo um ciclo de 41 000 anos (Berger & Loutre, 2004). O aumento da
12
F. Naughton, 2007
obliquidade provoca um aumento do contraste sazonal nas altas latitudes e
em particular, invernos muito frios e verões muito quentes em ambos os
hemisférios. Quando o valor da obliquidade decresce, o contraste sazonal
diminui em ambos os hemisférios e, a presença de verões amenos e de
invernos húmidos favorece o crescimento das calotes glaciárias nos pólos.
A precessão dos equinócios resulta da combinação de dois
movimentos de precessão: axial e de elipse (Fig. I.2.c). A precessão axial
resulta da modificação da orientação do eixo de rotação da terra
relativamente ao periélio e ao afélio, descrevendo uma figura cónica em
redor de uma recta perpendicular ao plano da eclíptica, a qual é
provocada pela força de atracção exercida pelo sol e lua à terra, a nível do
equador (Ruddiman, 2001). A precessão da elipse resulta do movimento da
rotação da terra sobre a órbita terrestre.
A precessão dos equinócios ocorre segundo ciclos de 23 000 a 19 000
anos (Ruddiman, 2001) sendo em média de 21 000 anos (Berger & Loutre,
2004). A variação da precessão dos equinócios produz um forte contraste
sazonal em ambos os hemisférios, onde verões quentes e invernos frios no
Hemisfério Norte contrastam com verões frios e invernos quentes no
Hemisfério Sul.
Fig. I.2 | Parâmetros astronómicos da terra: a) excentricidade, b) obliquidade e c) precessão dos
equinócios.
Nos últimos 450 000 anos, a alternância entre ciclos glaciários e
interglaciários foi dominada principalmente por variações da excentricidade
da órbita terrestre (Fig. I.1). As variações climáticas de longo-termo,
associadas à alternância entre os referidos ciclos glaciários e interglaciários,
foram detectadas pela primeira vez por Shackleton & Opdyke em 1973, no
registo de δ18O presente nas carapaças de foraminíferos bentónicos, numa
13
F. Naughton, 2007
sequência sedimentar marinha, confirmando pela primeira vez a teoria
astronómica. Posteriormente, esta variabilidade orbital foi observada noutros
registos marinhos, nomeadamente na sondagem ODP 980 (McManus et al.,
1999), assim como noutro tipo de registos tais como: no CO2 incluso nas
bolhas de ar da sondagem de gelo recolhida em Vostok, na Antárctica (Petit
et al., 1999) (Fig. I.3 B, C, D). A variabilidade orbital foi ainda detectada pela
primeira vez em sequências polínicas continentais nomeadamente, na
Grécia (Van der Hammen et al., 1971).
Por exemplo, a
sequência de
Tenaghi Philippon, evidência uma forte regressão da floresta durante os
períodos glaciários e uma forte expansão da mesma ao longo dos períodos
interglaciários (Fig. I.3 A). Existem ainda outros episódios de regressão da
floresta durante os períodos interglaciários, associados a ciclicidades
astronómicas de 20 000 a 40 000 anos (Fig. I.3 A, E).
Fig. I.3 | Variabilidade climática orbital nos registos marinhos, continentais e de gelo (adaptado de
Tzedakis et al., 2003). A: curva contínua a negro representa a percentagem de árvores excluindo
Juniperus e Pinus e a curva a tracejado representa a totalidade percentual de árvores; B: curva isotópica
do oxigénio contido nas carapaças de foraminíferos planctónicos e bentónicos da sequência ODP 980
(McManus et al., 1999); C: representa o volume de gelo versus nível do mar, obtido a partir do δ18O de
foraminíferos bentónicos da sequência ODP 980 (McManus et al., 1999); D: teor em CO2 atmosférico
contido na sondagem de gelo de Vostok (Petit et al., 1999); E: curva de insolação a 40°N e 65°N (Berger,
1978).
14
F. Naughton, 2007
1. 1. 2 Variabilidade climática milenar
1. 1. 2. 1 Os eventos de Heinrich e os eventos de D-O
Sobreposta à variabilidade climática orbital, ocorreram uma série de
flutuações rápidas, durante o último período glaciário (70 000-15 000 anos
calendário BP), cuja periodicidade não pode ser explicada pela teoria
orbital de Milankovitch. Esta variabilidade climática sub-orbital, ocorreu de
forma cíclica todos os 1470-1500 anos (Bond et al., 1993; Bond et al., 1997;
Mayeswski et al., 1997; Schulz et al., 2004) e é representada por uma
alternância entre episódios de aquecimento abrupto, designados por
interestadiais
e
episódios
de
arrefecimento
gradual
(estadiais), que
ocorreram ao longo do último período glaciário (Dansgaard et al., 1993). Esta
alternância entre episódios interestadiais (GIS-Greenland interstadials) e
estadiais (GS-Greenland stadials) é designada por oscilações de DansgaardOeschger (D-O). As oscilações de D-O foram detectadas pela primeira vez
no registo isotópico do oxigénio (δ18O) do gelo, na sondagem GRIP
(European Greenland Ice-core project) (Dansgaard et al., 1993) (Fig. I.4) (Fig.
I.5). A variação da temperatura associada a esta oscilação de D-O chegou
a atingir 16°C na Gronelândia (Severinghaus & Brook, 1999).
Fig. I.4 | Curva de variação da composição isotópica do oxigénio contido na sondagem de gelo GRIP
(Dansgaard et al., 1993). O valor Isotópico do oxigénio (δ18O) contido no gelo representa indirectamente a
temperatura atmosférica do momento no qual ocorreu a acumulação de gelo no pólo Norte (Johnsen et
al., 1992). Os números de 1 a 19 representam os episódios quentes interestadiais. Na parte superior da
figura estão representados os estádios isotópicos marinhos (MIS-Marine isotopic stage) os quais foram
definidos a partir da curva de variação do δ18O contido nas carapaças de foraminíferos bentónicos
(Shackleton & Opdyke, 1973). O MIS 3 terminou há cerca de: a) 24 000 anos segundo Shackleton &
Opdyke (1973) e Martinson et al. (1987) ou: b) 29 000 anos segundo Voelker et al. (1998) e van Kreveld et
al. (2000). Durante o último período glaciário, a variabilidade de D-O é mais proeminente durante o MIS 3
do que durante o MIS 2 (Voelker et al., 2002) e o MIS 4.
15
F. Naughton, 2007
Fig. I.5 | Localização de alguns dos registos paleoclimáticos citados no texto: GRIP (Dansgaard et al.,
1993); Vostok (Blunier et al., 1998; Petit et al., 1999); Byrd (Blunier et al., 1998); V23-81 (Bond et al., 1993;
Bond & Lotti, 1995); ENAM93-21 (Rasmussen et al., 1996; 1997); SU90-24 e SU90-16 (Elliot et al., 1998); SO825 (van Kreveld et al., 2000); NOAMP (Heinrich, 1988) (a sondagem marinha ODP609 localiza-se
exactamente na posição da NOAMP, Bond et al., 1993); HU75-55, 56 e HU87-09 (Andrews & Tedesco,
1992); MD95-2002 (Grousset et al., 2000); PS2644 (Voelker et al., 1998); ENAM97-09 (Richter et al., 2001);
SU90-38 (Cortijo et al., 1995); SU90-03 (Chapman & Shackleton,1998; Chapman & Maslin,1999; Chapman et
al., 2000); A (margem Ibérica): D11957P; SO75-26KL; PO 28-1; PO 8-2; MD95-2042; SU81-18; MD95-2041;
MD95-2040 e MD95-2039) (Lebreiro et al.,1996; Baas et al., 1997; 1998; Zahn et al., 1997; Abrantes et al.,
1998; Bard et al., 2000; Thouveny et al., 2000; de Abreu et al., 2003); RC11-83 (Charles et al., 1996); TTN05710/13/21 (Kanfoush et al., 2000). O rectângulo cinza claro representa a localização da cintura de
Ruddiman.
A
alternância
entre
fases
de
aquecimento
abrupto
e
de
arrefecimento gradual dos valores de temperatura da massa oceânica
superficial foi detectada ainda, em várias sondagens marinhas colhidas no
oceano Atlântico Norte, nomeadamente na V23-81 e ODP 609 (Fig. I.5)
(Bond et al., 1993; Bond & Lotti, 1995).
16
F. Naughton, 2007
Os episódios frios, ditos estadiais (GS), estão especialmente bem
representados nos registos obtidos em sondagens marinhas das altas
latitudes do Atlântico Norte, pela presença de níveis sedimentares pouco
espessos, ricos em material detrítico grosseiro, designado por IRD (Ice rafted
detritus), proveniente da descarga de icebergues (Bond & Lotti., 1995; Elliot
et al., 1998; van Kreveld et al., 2000).
Alguns destes episódios estadiais são considerados como episódios
extremos, e são conhecidos por eventos de Heinrich (H) (Broecker, 1994). Os
eventos de H ocorreram ciclicamente todos os 5 000-10 000 anos (Elliot et al.,
1998), e foram identificados pela primeira vez por Heinrich (1988), em
sondagens marinhas colhidas entre 45 e 50° N, na chamada “cintura de
Ruddiman” (Fig. I.5). Os níveis sedimentares representantes dos eventos de H
são caracterizados pela abundância anómala de IRDs (Fig. I.6) provenientes
das calotes glaciárias da “Laurentide”, Fenoescandinávia, Islândia e
Britânico-Irlandesa (Fig. I.7) (Heinrich, 1988; Andrews & Tedesco, 1992;
Broecker, 1994; Bond et al., 1992; Grousset et al., 1993, 2000; Bond & Lotti,
1995; Elliot et al., 1998; Hemming et al., 1998; Scourse et al., 2000; Richter et al.,
2001), assim como por um aumento dos valores de susceptibilidade
magnética (MS) (Grousset et al., 1993). A presença de IRDs foi detectada,
ainda, fora da cintura de Ruddiman, a norte de 50°N (Cortijo et al., 1995 ;
Fronval et al., 1995; Rasmussen et al., 1996; 1997; Revel et al., 1996 ; Andrews
et al., 1998; Elliot et al., 1998; Voelker et al., 1998; Van Kreveld et al., 2000),
assim como nas médias latitudes do Atlântico Norte, a sul de 45° N (Lebreiro
et al.,1996; Baas et al., 1997; 1998; Zahn et al., 1997; Abrantes et al., 1998; Bard
et al., 2000; Thouveny et al., 2000; de Abreu et al., 2003; Chapman &
Shackleton, 1998; Chapman & Maslin, 1999; Chapman et al., 2000) (Fig. I.5). A
espessura das camadas de IRDs (IRD-layers) detectada nas sondagens
marinhas das médias latitudes é, contudo, mais fina do que a espessura das
camadas que caracterizam os eventos de Heinrich ao longo da cintura de
Ruddiman. O mesmo se passa relativamente aos valores de MS, durante um
evento de Heinrich: estes são elevados na cintura de Ruddiman e
relativamente baixos nas latitudes médias do Atlântico Norte (Thouveny et
al., 2000).
17
F. Naughton, 2007
Fig. I.6 | Identificação dos eventos de Heinrich na sondagem ODP609 (Bond et al., 1993). Curva de
variação: da % de foraminíferos planctónicos de origem polar (N. pachyderma sin.); de detritos
provenientes das calotes glaciárias do Hemisfério Norte (IRD); de δ18O contido nas carapaças de N.
pachyderma sin. e cujos picos representam uma grande quantidade de fluxo de água fundida (Hemming,
2004).
Fig. I.7 | Calotes glaciárias (Ruddiman, 2001).
18
F. Naughton, 2007
Apesar da espessura das camadas de IRDs e do sinal magnético
serem menos importantes nas médias do que nas altas latitudes do Atlântico
Norte, o impacto destes eventos extremos é evidente, principalmente no que
se refere à diminuição brutal dos valores da temperatura da massa de água
superficial (SST-Sea surface temperature) em ambas as regiões. Este drástico
arrefecimento da massa de água superficial, resultou da introdução de
grandes quantidades de água, proveniente da fusão dos icebergues no
Atlântico Norte (Rosell-Melé et al., 1997; Bard et al., 2000; van Kreveld et al.,
2000; Pailler & Bard, 2002). A diminuição da SST é particularmente bem
representada no sinal isotópico do oxigénio incluso nas carapaças de
foraminíferos planctónicos do tipo Globigerina bulloides (Maslin et al., 1995;
Bond & Lotti, 1995; Bond et al., 1993; Cortijo et al., 1997; Chapman &
Shackleton, 1998; Chapman & Maslin, 1999; Chapman et al., 2000;
Shackleton et al., 2000a; Schönfeld et al., 2003), assim como, no aumento da
abundância
de
foraminíferos
planctónicos
polares
do
tipo
Neogloboquadrina pachyderma sinistrógira (Bond et al., 1992; Fronval et al.,
1995; Cortijo et al., 1997; Lebreiro et al., 1997; Rasmussen et al., 1997; Cayre et
al., 1999; de Abreu et al., 2003) (Fig. I.6).
Para além da diminuição drástica da SST, a introdução de grandes
quantidades de água doce, provocou uma diminuição dos valores de
salinidade na massa de água oceânica de superfície (SSS-Sea surface
salinity) (Vidal et al., 1997; 1999), produzindo mudanças paleoceanográficas
muito importantes (Lehman & Keigwin, 1992; Maslin et al., 1995; Rahmstorf,
1995; Cortijo et al., 1997; Vidal et al., 1997; Zahn et al., 1997; Chapman &
Schackelton, 1998). A variabilidade dos parâmetros SST e SSS observada
durante os eventos de Heinrich induziu a uma forte estratificação da coluna
de água, afectando o padrão natural da
circulação
termohalina
(circulação global oceânica, também conhecida por “conveyor belt”;
Broecker, 1991) (Fig. I.8) e, em particular, da “Atlantic Meridional Overturning
Circulation” (MOC) impedindo a transferência de calor das baixas para as
altas latitudes, provocando um forte arrefecimento do Atlântico Norte
(Fawcett et al., 1997).
Para além da redução ou interrupção da MOC, a introdução de
grandes quantidades de água doce impediu, chegando mesmo a bloquear
19
F. Naughton, 2007
totalmente, a formação da massa de água profunda, no Atlântico Norte
(NADW-North Atlantic Deep Water) (Broecker, 1994; Keigwin & Lehman, 1994;
Maslin et al., 1995; Oppo & Lehman, 1995; Vidal et al., 1997; Zahn et al., 1997;
Seidov & Maslin, 1999; Ganopolski & Rahmstorf, 2001; Elliot et al., 2002). A
formação da NADW pode ser ainda perturbada por variações na posição
da frente polar, a qual impede o transporte da corrente superficial Norte
Atlântica para as altas latitudes do Atlântico Norte, o consequente
arrefecimento da massa de água superficial, o mergulho da mesma para o
oceano profundo e finalmente o retorno para sul dessa massa de água
profunda (Broecker et al., 1990).
Fig. I.8 | Representação esquemática do padrão geral da circulação termohalina (Rahmstorf, 2002).
MOC- Atlantic Meridional Overturning Circulation.
Nas últimas duas décadas, vários mecanismos têm sido propostos
para explicar a ocorrência dos eventos de Heinrich (ver Hemming, 2004) no
Atlântico Norte, nomeadamente:
1- Instabilidade interna da calote glaciar da “Laurentide” (“bingepurge model”, MacAyeal, 1993; Verbitsky & Saltzaman, 1995; Clark et al.,
1996; Papa et al., 2006). Este mecanismo pressupõe que o calor geotermal é
o principal responsável pela descarga de icebergues no Atlântico Norte;
2- Actividade repetitiva de fenómenos de “jökulhaup” no lago da Baía
de Hudson (Johnson & Lauritzen, 1995). Esta actividade jökulhaups representa
20
F. Naughton, 2007
a ocorrência de sucessivos episódios drásticos de expulsão de água os quais
estão directamente relacionados com variações do nível do lago;
3- Crescimento de plataformas de gelo e consequente colapso no
Mar de Labrador (Hulbe, 1997; Hulbe et al., 2004; Alley et al., 2006).
A diminuição da SST, detectada nos vários registos do Atlântico Norte,
durante os episódios estadiais de D-O é de menor amplitude do que aquela
observada durante os episódios estadiais extremos (eventos de Heinrich).
Contudo, estes arrefecimentos são, tal como durante os eventos de Heinrich,
principalmente induzidos pela introdução de grandes quantidades de água
doce no Atlântico Norte, as quais perturbaram a circulação termohalina
durante o último período glaciário (van Kreveld et al., 2000; Boyle, 2000;
Ganopolski & Rahmstorf, 2001; Elliot et al., 2002; Knutti et al., 2004). Para além
da alteração no modo de funcionamento da circulação termohalina (THC),
a formação da NADW foi igualmente afectada: sendo esta reduzida durante
os episódios estadiais e semelhante à situação actual, durante os episódios
interestadiais.
O efeito produzido por tais modificações hidrológicas, que ocorreram
no Atlântico Norte, foi transmitido globalmente de forma súbita e está bem
representado em varios registos paleoclimáticos (Leuschner & Siroko, 2000;
ver Voelker et al., 2002) (Fig. I.9). No entanto, esta interpretação associada a
uma reorganização rápida entre o oceano e a atmosfera foi recentemente
posta em questão por Wunsch (2006). Este autor propõe que mudanças do
volume de gelo no pólo Norte teram sido responsáveis por variações da
direcção dos ventos as quais afectaram por consequência o oceano.
21
F. Naughton, 2007
Fig. I.9 | Distribuição global dos registos de D-O durante o MIS 3 (ver Voelker et al., 2002).
Algumas zonas, tais como o Oceano Atlântico Sul (Charles et al., 1996;
Vidal et al., 1999; Kanfoush et al., 2000) e a Antárctica (cores de gelo Byrd e
Vostok) (Blunier et al., 1998), apresentam um registo climático de ordem
temporal milenar em anti-fase, quando comparado com os registos obtidos
na Gronelândia e Atlântico Norte (Bender et al., 1994; Blunier et al., 1998;
Blunier & Brook, 2001) (Fig. I.5). Para além deste registo em anti-fase, a
variabilidade climática observada no Hemisfério Sul é bastante mais fraca
em amplitude do que aquela registada na Gronelândia.
A correlação anti-fásica entre os dois hemisférios é explicada pelo
designado efeito de “thermal bipolar sea-saw”, o qual sugere que variações
abruptas na intensidade da THC, causadas por modificações no “input” de
água doce, afecta o clima nos pólos através de modificações associadas à
transferência de calor meridional (Ganopolski & Rahmstorf, 2001; Knutti et al.,
2004).
Este efeito de “sea-saw” é um dos muitos mecanismos propostos para
explicar a passagem abrupta de um evento frio (estadial ou de Heinrich) a
um evento interestadial de D-O, no Hemisfério Norte (Knutti et al., 2004). Para
além desta teoria, vários têm sido os mecanismos propostos na tentativa de
explicar este fenómeno, nomeadamente:
22
F. Naughton, 2007
- concentração de sal durante a formação de gelo marinho nas zonas
de convecção do Atlântico Norte (van Kreveld et al., 2000). Durante a
acumulação de gelo marinho, o sal é rejeitado e introduzido no oceano
provocando o aumento da intensidade da MOC (Atlantic Meridional
Overturning). Este mecanismo foi confirmado recentemente através de um
modelo numérico de complexidade intermédia (Earth system Model of
Intermediate Complexity) efectuado por Wang et al. (2006).
- “feedback” oceânico o qual envolve a corrente oceânica
mediterrânea (MOW-Mediterranean Outflow Water) (Voelker et al., 2006).
Este mecanismo sugere que o aumento na intensidade da MOW que
ocorreu durante episódios de fraca intensidade da circulação termohalina
(THC) levou à introdução de calor e sal, proveniente da MOW, nas massas
de água intermédias durante os episódios frios o que terá levado a THC a
mudar o seu modo de funcionamento;
- mudanças de sazonalidade, nomeadamente variações severas na
temperatura de inverno as quais são provocadas por variações na extensão
de gelo marinho no Atlântico Norte (Denton et al., 2005). No inverno, durante
a formação de grandes camadas de gelo marinho, o sal incorporado na
água do mar é rejeitado e a sua incorporação nas massas de água
intermédias favorece o restabelecimento da MOC.
Para além desta problemática associada à passagem abrupta de
condições estadiais a condições interestadiais, outras questões têm sido
levantadas nomeadamente em relação ao tipo de mecanismos que agiram
de forma a amplificar o sinal climático, numa dada região, durante os
eventos de Heinrich, nomeadamente:
- a subida do nível do mar, causada pela introdução de grandes
quantidades de água doce no Atlântico Norte, durante os eventos de
Heinrich, contribuíram para a destabilização das plataformas de gelo e das
calotes glaciárias do Hemisfério Norte induzindo uma contínua descarga de
23
F. Naughton, 2007
icebergs amplificando o sinal destes eventos nas altas latitudes do Atlântico
Norte (Flückiger et al., 2006);
- mecanismo atmosférico: a presença de um elevado índice na
oscilação Norte Atlântica (NAO-North Atlantic Oscillation) provocaria a
migração dos centros de alta pressão móveis polares Escandinavo e
Atlântico, para a Península Ibérica, contribuindo para uma amplificação das
condições extremamente frias e áridas, durante os eventos de Heinrich, nas
regiões situadas mais a oeste da zona mediterrânica as quais não têm
ligação directa com o Atlântico Norte (Sánchez Goñi et al., 2002).
Todos estes mecanismos citados anteriormente sugerem que as
oscilações de D-O resultam de modificações que ocorreram inicialmente no
Atlântico Norte e cuja resultante foi posteriormente transmitida globalmente.
Contudo, nos últimos anos, outras teorias têm sido enunciadas na tentativa
de explicar a variabilidade de D-O, as quais sugerem que esta variabilidade
global envolveu outro tipo de mecanismos que ocorreram inicialmente na
zona tropical, tais como: monções indianas e asiáticas (Leuschner & Sirocco,
2000; Kudrass et al., 2001) e as modificações que ocorreram na zona do
Pacífico equatorial associadas ao El Niño Southern Oscilation (ENSO) (Cane &
Clement, 1999).
Recentemente, Wunsch (2006), sugeriu ainda uma alternativa à teoria
do “freshwater pulse mechanism” (associada à introdução de grandes
quantidades de água no oceano), a qual sugere que as modificações da
circulação oceânica resultam de variações na intensidade dos ventos e
que, os eventos de D-O resultam de uma interacção entre o vento e a
topografia das calotes glaciárias permitindo assim a transferência rápida do
sinal climático a nível global.
Contudo, até à data, os vários mecanismos propostos na tentativa de
explicar a natureza quase cíclica dos eventos de D-O e de Heinrich não são
conclusivos.
24
F. Naughton, 2007
Para além da impressão digital deixada por esta variabilidade
climática nas sondagens de gelo da Gronelândia e nas sondagens marinhas
do Atlântico Norte, estes eventos tiveram também um forte impacto nas
variações do coberto vegetal, nomeadamente nas médias latitudes, uma
vez que no passado esta zona nunca terá sido coberta por gelo.
A correlação efectuada entre os vários registos continentais e os
registos paleoclimáticos obtidos tanto em sequências marinhas como em
sondagens de gelo não é directa, uma vez que cada um destes locais
apresenta um modelo cronológico próprio. Por estas razões, o estudo de
sondagens marinhas profundas ricas em conteúdo polínico, situadas nas
proximidades do continente, é muito importante na compreensão da
resposta da vegetação à variabilidade climática detectada no Atlântico
Norte, uma vez que permite correlacionar directamente ambos os registos
marinho e continental utilizando uma cronologia única (Groot & Groot, 1966;
Balsam & Heusser, 1976; Heusser & Shackleton, 1979).
De facto, a correlação directa dos diferentes tipos de registos
paleoclimáticos permite-nos não só identificar a variabilidade climática que
afectou
o
continente
mas
também
correlacioná-la
com
os
outros
componentes do sistema climático (oceano-atmosfera-gelo).
A margem ibérica (situada a latitudes médias do Atlântico Norte) é
directamente influenciada pela variabilidade climática detectada no
Atlântico Norte e Gronelândia, sendo por isso considerada uma zona
preferencial para efectuar este tipo de correlação directa.
Até à data, foram efectuadas várias correlações directas (oceanocontinente ou oceano-continente-gelo) ao longo da margem ibérica, na
tentativa de compreender como é que a vegetação respondeu à
variabilidade climática que caracteriza o último período glaciário (o qual
engloba os estádios isotópicos marinhos MIS 4, MIS 3 e MIS 2), nas latitudes
médias do Atlântico nordeste, nomeadamente: durante o MIS 3 Combourieu Nebout et al., 1999, 2002; Sánchez Goñi et al., 2000, 2002;
Roucoux et al., 2001, 2005, e durante o MIS 2 - Hooghiemstra et al., 1992;
Combourieu Nebout et al., 1999, 2002; Boessenkool et al., 2001; Roucoux et
al., 2001, 2005; Turon et al., 2003 (Fig. I.10). Contudo, a resolução temporal
25
F. Naughton, 2007
utilizada para o estudo do MIS 2 nas várias sequências sedimentares é
bastante baixa.
Fig. I.10 | Localização das sondagens onde foi efectuado uma correlação directa oceano-continente:
8057B (Hooghiemstra et al., 1992), ODP 976 (Combourieu Nebout, et al., 1999; 2002), MD95-2042 (Sánchez
Goñi et al., 2000; 2002), SO75-6KL (Boessenkool et al., 2001), MD95-2039 (Roucoux et al., 2001; 2005), MD952043 (Sánchez Goñi et al., 2002) e SU81-18 (Turon et al., 2003). Delimitação das zonas biogeográficas
(adaptado de Peinado & Rivas-Martinez, 1987).
Para além da baixa resolução temporal, a maioria dos estudos foram
efectuados em áreas adjacentes à zona biogeográfica mediterrânica
(Hooghiemstra et al., 1992; Combourieu Nebout et al., 1999, 2002;
Boessenkool et al., 2001; Turon et al., 2003) a qual é caracterizada por um
período de seca estival (Fig. I.10). A única correlação directa oceanocontinente-gelo elaborada nas proximidades da zona biogeográfica
Eurosiberiana (numa zona de transição entre esta e a zona biogeográfica
Mediterrânica), caracterizada por um clima mais húmido, foi efectuada por
Roucoux et al. (2005), (Fig. I.10).
A maioria destas correlações sugere que a vegetação respondeu de
forma síncrona à variabilidade da SST ou seja, a diminuição da temperatura
da massa de água superficial ocorreu contemporaneamente à diminuição
26
F. Naughton, 2007
da floresta decídua/mediterrânica e à forte expansão da vegetação semidesértica no continente, enquanto que, durante os episódios quentes,
ocorreu uma forte expansão da floresta mediterrânica e uma forte redução
da vegetação semi-desértica.
Alguns destes trabalhos (Boessenkool et al., 2001; Roucoux et al., 2005)
delimitam os eventos de Heinrich baseados apenas na espessura da
camada de IRD, omitindo o registo dos indicadores de SST o qual permite
definir o intervalo completo destes episódios.
Outros estudos, efectuados ao longo da margem Ibérica, detectaram
um
padrão
sedimentar
complexo
durante
os
eventos
de
Heinrich
nomeadamente, nas sondagens marinhas: PO 28-1 e a PO 8-2 (Abrantes et
al., 1998), MD95-2039 (Thouveny et al., 2000; Schönfeld et al., 2003), MD952042 (Sánchez Goñi et al., 2000; Thouveny et al., 2000), SU81-18 (Bard et al.,
2000), MD95-2040 (Schönfeld et al., 2003; Narciso et al., 2006).
A conjugação dos dois indicadores de presença de icebergues ao
longo da margem ibérica, IRD e susceptibilidade magnética, efectuada nas
sondagens marinhas SU81-18, MD95-2039, MD95-2042, permitiu detectar a
ocorrência de dois sub-episódios associados aos eventos de H2 e H1 (Bard et
al., 2000; Thouveny et al., 2000). O sub-episódio mais antigo apresenta uma
fraca quantidade de IRDs e um grande pico no sinal magnético, enquanto
que o mais recente, é representado por uma grande quantidade de IRDs e
valores elevados em MS.
Contemporaneamente aos episódios complexos de deposição de
IRDs, ocorreu uma forte diminuição da SST, sugerindo que, apesar da
ausência deste material grosseiro, o impacto dos eventos de Heinrich é bem
evidente ao longo da margem Ibérica (Lebreiro et al., 1997; Cayre et al.,
1999; Bard et al., 2000; Pailler & Bard, 2002; Chapman et al., 2000; Shackleton
et al., 2000a; de Abreu et al., 2003; Schönfeld et al., 2003). Apesar dos
indicadores marinhos apresentarem um padrão complexo relativamente aos
eventos de H2 e H1, na sondagem SU81-18, o continente adjacente é
caracterizado por um sinal homogéneo representante de condições áridas
(Turon et al., 2003). No entanto, a correlação directa continente-oceano de
alta resolução temporal, efectuada na sondagem MD95-2042 (situada nas
proximidades da SU81-18), sugere a presença de um padrão tri-fásico
27
F. Naughton, 2007
associado aos eventos de Heinrich durante o MIS 3 (Sánchez Goñi et al.,
2000), evidente no sinal obtido tanto a partir dos indicadores marinhos como
dos indicadores continentais. Cada um dos eventos de H5, H4 e H3 é
inicialmente marcado por condições húmidas no continente, seguido por
um aumento de aridez durante o episódio de máxima deposição de IRDs
(provenientes das calotes glaciárias Canadianas) e finalmente, por uma
nova fase húmida no final de cada um desses eventos.
Várias hipóteses foram enunciadas na tentativa de explicar a
ocorrência deste padrão complexo ao longo da margem Ibérica, durante os
eventos de Heinrich, nomeadamente:
a) a proveniência de IRD resulta de diferentes fontes (Bard et al., 2000,
Thouveny et al., 2000);
b) resulta de múltiplas descargas de icebergues das calotes glaciárias
canadianas (Abrantes et al., 1998); ou
c) migração da frente polar (Chapman et al., 2000).
Contudo, nenhum destes mecanismos propostos permite explicar
variações entre humidade e aridez detectadas no continente.
Objectivo 1
Por estas razões, um dos principais objectivos deste trabalho é,
identificar e compreender a resposta da vegetação ao complexo sinal
deixado pelos eventos de Heinrich, no noroeste da margem Ibérica, assim
como de discutir sobre os eventuais mecanismos responsáveis por esta
variabilidade, através da comparação entre dados e modelos numéricos.
Especial atenção será dada aos eventos de Heinrich que ocorreram durante
o MIS 2 (26 000-15 500 anos cal BP) (intervalo cronológico definido para o MIS
2 por Shackleton & Opdyke. 1973), nomeadamente H2 e H1. De forma a
alcançar estes objectivos, foi efectuada uma correlação directa (oceanocontinente-gelo) de altíssima resolução temporal (< 200 anos) numa região
directamente influenciada pelo clima das latitudes temperadas que se
encontre fora da zona sub-tropical (Margem Ibérica).
28
F. Naughton, 2007
Os episódios de H2 e H1, são separados por um período associado a
condições glaciárias plenas (máxima extensão de gelo nos pólos), entre
24 300 e 18 500 anos cal BP, designado por último máximo glaciar (LGM-Last
Glacial Maximum). Apesar do nível do mar ter atingido o seu valor mínimo,
entre 30 000 e 20 000 anos cal BP, o LGM foi definido preferencialmente num
período caracterizado por uma certa estabilidade climática, na qual não
existem variações bruscas tais como eventos de Heinrich e de D-O (Mix et al.,
2001).
O LGM é considerado como um período chave na compreensão da
sensibilidade a mudanças dos vários sistemas ambientais globais uma vez
que o clima permaneceu relativamente estável, embora bastante diferente
e oposto ao estado actual (condições interglaciárias) (Mix et al., 2001).
Recentemente foram efectuadas, no âmbito do projecto MARGO
(Multiproxy Approach for the Reconstruction of the Glacial Ocean surface),
uma série de reconstruções das condições oceânicas superficiais (SST e
extensão de gelo marinho) para o LGM, nomeadamente na zona do
Atlântico e Mares do Norte. Contrariamente às reconstruções prévias,
obtidas no âmbito do projecto CLIMAP (Climate Long-range Investigation,
Mapping, And Prediction) (CLIMAP project members, 1981), as quais
sugeriam a presença de uma grande extensão de coberto de gelo perene a
norte do Atlântico Norte e Mares do Norte durante este período, o grupo
Margo mostra que esta cobertura de gelo de mar seria bastante mais
variável (Kucera et al., 2005; Meland et al., 2005). Esta variabilidade resulta
essencialmente do forte contraste sazonal que caracteriza este período, ou
seja: invernos frios que favorecem uma forte expansão de gelo marinho
enquanto que verões quentes reduzem fortemente a área coberta por gelo
no Atlântico e Mares do Norte (de Vernal et al., 2005a). Esta diferença entre
condições perenes e sazonais tem implicações muito importantes no clima,
nomeadamente no que se refere à quantidade de humidade fornecida às
altas latitudes do Hemisfério Norte assim como na localização da zona
preferencial de formação da NADW e ainda, na Intensidade e extensão da
MOC (Meland et al., 2005).
29
F. Naughton, 2007
Nos últimos anos foram efectuadas várias reconstruções de SST ao
longo da margem Ibérica as quais incluem o LGM. Estas reconstruções foram
essencialmente
baseadas
na
composição
química
de
alcanonas
“alkenones” (Bard et al., 2000; Pailler & Bard, 2002) e na aplicação de
funções de transferência às associações de foraminíferos planctónicos
(Lebreiro et al., 1997; Cayre et al., 1999; de Abreu et al., 2003). Os resultados
obtidos nessas reconstituições mostram que a temperatura média anual seria
igual ou superior a 13°C; a de verão seria igual ou superior a 15°C e a
temperatura de inverno situada entre 12 a 14°C, ao longo do LGM.
Para além do aquecimento da massa de água superficial durante o
LGM, algumas das reconstruções obtidas pelo grupo MARGO, mostram uma
gradual diminuição da SST anual segundo um perfil latitudinal de sul a norte
nomeadamente: cerca de 18°C na extremidade sudoeste Portuguesa, de
16°C à latitude de Lisboa, de 14°C à latitude do Porto e finalmente de 12°C
na extremidade noroeste da Península Ibérica (Morey et al., 2005).
Baseado no trabalho desenvolvido por de Van Campo (1984),
Sánchez Goñi (2006) sugeriu que o desenvolvimento da floresta temperada,
durante os interestadiais de D-O que ocorreram ao longo do MIS 3, está
intimamente associada a valores de SST de verão iguais ou superiores a 12°C.
De facto, o aumento da SST de verão durante o LGM favoreceu o
restabelecimento da “Meridional Overturnig Circulation” (MOC) permitindo a
transferência de calor e humidade para o continente europeu, incluindo a
margem Ibérica.
Desta forma, e como resposta à variação da SST de verão estimada
pelos trabalhos referidos anteriormente, onde os valores da SST de verão são
da ordem dos 15°C, deveríamos esperar encontrar uma forte expansão da
floresta temperada no noroeste da Península Ibérica.
No entanto, o diagrama polínico obtido para a sondagem MD95-2039
(Roucoux et al., 2005), mostra uma maior expansão de árvores temperadas
durante os episódios interestadiais de D-O especialmente entre 65 000 e 30
000 anos cal BP do que durante o LGM, apesar dos valores de SST serem
semelhantes ao longo dos referidos episódios.
30
F. Naughton, 2007
A maioria dos trabalhos de correlação directa (oceano-continente e
oceano-continente-gelo) efectuados ao longo da margem ibérica tem
dado pouca importância ao período que caracteriza o último máximo
glaciar (LGM).
Objectivo 2
Desta forma, pretende-se identificar a resposta da vegetação ao
longo do período de máxima extensão de gelo nos pólos e relacioná-la com
as variações das condições oceânicas de superfície que caracterizam o
LGM. Para melhor compreender a relação entre a redução da floresta
temperada do noroeste da Península Ibérica e as condições de SST da
margem adjacente pretende-se ainda comparar os dados obtidos para o
LGM com outros referentes ao interestadiais de D-O do tardi-MIS 3.
Ambiciona-se ainda discutir sobre potenciais mecanismos responsáveis pela
amplificação ou redução do sinal climático continental, nomeadamente o
impacto do volume de gelo acumulado nos pólos e ainda as variações
sazonais de expansão de gelo marinho nas altas latitudes do Atlântico Norte,
nas latitudes médias.
A variabilidade climática milenar, para além de ter sido detectada ao
longo do último período glaciário (MIS 4, MIS 3 e MIS 2), foi ainda observada
durante o MIS 1, o qual engloba o período de transição entre o último
episódio glaciário e o actual interglaciário (LGIT-Last Glacial Interglacial
transition) (15 500 – 11 500 cal anos BP) e o Holocénico (últimos 11 500 cal
anos BP).
1. 1. 2. 2 O início da deglaciação
A fase de transição entre o último período glaciário e o actual
interglaciário (LGIT: Last Glacial-Interglacial Transition) é caracterizada por
um valor máximo da insolação de verão a 65° N e por uma forte redução do
volume de gelo acumulado nos pólos. Apesar das condições orbitais
favorecerem o início de um episódio interglaciário durante o LGIT, este
período é, contudo marcado por uma série de oscilações climáticas
complexas e abruptas.
31
F. Naughton, 2007
Esta variabilidade climática foi detectada pela primeira vez numa
série de sequências polínicas continentais norte europeias e, os seus
sucessivos eventos foram designados por: Oldest Dryas (frio), Bølling (quente),
Older Dryas (frio), Allerød (quente) e Younger Dryas (frio) (Iversen, 1954;
Mangerud et al., 1974). Na Península Ibérica, a qual se situa nas latitudes
médias do Atlântico Norte, apenas três destes episódios foram detectados: o
Oldest Dryas (Dryas Antigo), o Bølling-Allerød e o Younger Dryas (Dryas
Recente) (Pons & Reille, 1988; de Beaulieu et al., 1994; Peñalba et al., 1997;
Von Engelbrechten, 1998).
Posteriormente, a variabilidade climática associada ao LGIT foi
observada nas sondagens de gelo da Gronelândia e os seus eventos
designados por: GS 2 (Greenland stadial 2); GIS 1 (Greenland interstadial 1) e
GS 1 (Greenland stadial 1) (Alley et al., 1993; Dansgaard et al., 1993; Johnsen
et al., 2001) e ainda numa série de sequências marinhas recolhidas no
Atlântico Norte (Ruddiman & McIntyre, 1981; Lehman & Keigwin, 1992; Bond
et al., 1993; Rasmussen et al., 1996). Estes registos marinhos do Atlântico Norte
mostram um aumento abrupto na temperatura da massa de água
superficial após o evento de H1 (Severinghaus & Brook, 1999), o qual
permaneceu mais ou menos estável durante 2000 anos (Broecker, 2000). Este
fenómeno está intimamente correlacionado com o evento “meltwater pulse
1A” detectado no registo dos corais de Barbados (Bard et al., 1990a). O
episódio quente oceânico associado ao evento continental Bølling-Allerød
(B-A), foi seguido por um arrefecimento rápido e intenso (Ruddiman &
McIntyre, 1981; Lehman & Keigwin, 1992; Bond et al., 1993; Rasmussen et al.,
1996) que caracteriza o Dryas recente (Younger Dryas) no oceano. O Dryas
recente é considerado como o incidente climático mais abrupto que
ocorreu após o início da deglaciação (Broecker, 1994; 2000). O retorno a
condições glaciárias plenas ocorreu em simultaneo com a diminuição das
calotes glaciárias do Hemisfério Norte e, consequente introdução de
grandes quantidades de água doce no oceano, a qual perturbou o padrão
geral da circulação termohalina (Teller et al., 2002) assim como a formação
de NADW.
A mudança da direcção do fluxo de água doce, proveniente das
fases finais da fusão da calote glaciar da “Laurentide”, do rio Mississippi para
32
F. Naughton, 2007
o rio St. Laurence, é tida como o mecanismo principal responsável para a
ocorrência do Dryas recente no Atlântico Norte (Clark et al., 2001). Para
além da introdução de água doce proveniente da calote glaciar
canadiana no Atlântico Norte, a incorporação de água doce do lago de
gelo Báltico na Escandinávia terá contribuído também para a destabilização
da MOC durante este período (Nesje et al., 2004).
Para além do Atlântico Norte e continente Europeu, o Dryas recente
deixou indícios da sua presença nos mais variadíssimos registos climáticos
mundiais, nomeadamente: na América do Norte (Mott et al., 1986; Peteet et
al., 1990; 1993; Levesque et al., 1993; Peteet & Man, 1994), na América
Central e Caraíbas (Hughen et al., 1996; 2000; Peterson., 2000), na América
do Sul (van der Hammen & Hooghiemstra, 1995), em Africa (deMenocal.,
2000), no Mar Sulu situado no Pacífico tropical (Linsley & Thunell, 1990;
Rosenthal et al., 2003), no Pacífico noroeste (Kotilainen & Shackleton, 1995), a
norte do Mar Arábico (Schulz et al., 1998) e ainda no sudeste Atlântico
(situado a este da Bacia Angolana ~ a 17° S) (Kim et al., 2002) (Fig. I.11).
Fig. I.11 | Alguns registos da variabilidade climática milenar durante o LGIT (adaptado de Goslar et al.,
2000): Lago de Gosciaz na Polónia (Goslar et al., 1993); Bacia de Caríaco (Hughen et al., 1996); GRIP δ18O
(Dansgaard et al., 1993).
33
F. Naughton, 2007
Contudo e tal como acontece durante os eventos estadiais de D-O, o
arrefecimento da Antártida (ACR-Antarctic Cold Reversal) e das altas
latitudes do Hemisfério Sul ocorreu anteriormente ao arrefecimento que
caracteriza o Dryas recente na zona tropical e Atlântico Norte (Bender et al.,
1994; Sowers & Bender, 1995; Blunier et al., 1998; Blunier & Brook, 2001; Bianchi
& Gersonde, 2004) (Fig. I.12). Este sinal anti-fásico é, tal como para os eventos
de D-O, explicado pela teoria do “sea-saw”.
Fig. I.12 | Comparação dos registos de GRIP δ18O (Gronelândia) (Dansgaard et al., 1993) com δ18O das
sondagens gelo da Antárctica: Taylor Dome e Byrd e com a variação de Deutério na sondagem de gelo
Vostok (adaptado de Blunier et al., 1998). Esta correlação foi efectuada utilizando a curva de metano de
Vostok.
Durante a última deglaciação, a fusão terminal dos glaciares
continentais favoreceu a formação de numerosos lagos. Desta forma, os
sedimentos lacustres foram considerados, durante muito tempo, como locais
propícios ao estudo da variabilidade climática que caracteriza este período.
Infelizmente as sequências lacustres raramente apresentam registos mais
antigos que 15 000 anos.
Ao longo de vários anos foram efectuados uma série de estudos
polínicos em sequências lacustres de forma a detectar modificações no
coberto vegetal durante este período e, em particular, na Península Ibérica
(Fig. I.13).
34
F. Naughton, 2007
Fig. I.13 | Localização geográfica de alguns dos registos polínicos continentais. a) quadrado A localiza as
sequências de 1 a 5: 1- Laguna de la Roya (Allen et al., 1996), 2- Sanabria March (Allen et al., 1996), 3Laguna de las Sanguijuelas (Muñoz Sobrino et al., 2004), 4-Lleguna (Muñoz Sobrino et al., 2004), 5- Pozo do
Carballal (Muñoz Sobrino et al., 1997); b) os pontos 6 a 13 correspondem a: 6- Laguna Lucenza (Santos et
al., 2000); 7- Lagoa Lucenza (Muñoz Sobrino et al., 2001); 8-Lago de Ajo (Allen et al., 1996); 9- Los Tornos
(Peñalba, 1994); 10-Saldropo (Peñalba, 1994); 11-Belate (Peñalba, 1994); 12- Atxuri (Peñalba, 1994); 13Banyoles (Pérez-Obiol & Julià, 1994); c) quadrado B inclui as sequências 14 a 19: 14- Quintanar de la Sierra
(Peñalba et al., 1997); 15- Sierra de Neila - Quintanar de la Sierra (Ruiz Zapata et al., 2002); 16- Hoyos de
Iregua (Gil-Garcia et al., 2002); 17- Laguna Masegosa (Von Engelbrechten, 1998); 18- Laguna Negra (Von
Engelbrechten, 1998); 19- Las Pardillas lake (Sánchez Goñi & Hannon, 1999); 20- Padul (Pons & Reille, 1988);
21- Mougás (Gómez-Orellana et al., 1998); 22- Charco da Candieira (Van der Knaap & Van Leeuwen,
1995).
Contudo, e tal como foi referido anteriormente, a correlação
efectuada entre os vários registos continentais e os registos paleoclimáticos
obtidos tanto em sequências marinhas como em sondagens de gelo, é
indirecta. Para além disso, algumas sequências continentais apresentam por
vezes hiatos sedimentares importantes, os quais impedem a realização de
um estudo detalhado sobre a deglaciação, sendo por isso, necessário
recorrer a sondagens marinhas situadas nas proximidades do continente.
Até à data, foram efectuados apenas dois estudos de correlação
directa (oceano-continente) ao longo da margem ibérica, para o MIS 1
35
F. Naughton, 2007
(Boessenkool et al., 2001; Turon et al., 2003). Contudo, os resultados obtidos
não foram discutidos em pormenor para este período de transição.
Objectivo 3
Por estas razões pretende-se detectar a variabilidade climática que
ocorreu durante o LGIT no noroeste da Península Ibérica, nomeadamente:
documentar as modificações do coberto vegetal contemporâneas da
variabilidade climática detectada no Atlântico Norte e Gronelândia: fase
final do evento de Heinrich 1, o Interestadial GIS 1 (episódio quente B-A),
estádio GS 1 (Dryas recente) e o início do Holocénico.
1. 1. 2. 3 Holocénico
Durante muitos anos o actual interglaciário foi considerado como um
período onde o clima permaneceu relativamente estável.
No entanto, nos últimos anos, vários estudos efectuados em registos
marinhos, continentais e de gelo, assim como outros dados de modelos
numéricos, mostraram que o Holocénico foi afectado por uma variabilidade
climática de longo termo induzida por modificações orbitais (Kutzbach &
Gallimore, 1988; Koç & Jansen, 1994; Duplessy et al., 2001; Crucifix et al., 2002;
Marchal et al., 2002; Weber & Oerlemans, 2003; Andersen et al., 2004a; Moros
et al., 2004; Solignac et al., 2004; Keigwin et al., 2005; Renssen et al., 2005;
Lorenz
et
al.,
2006)
e
por
uma
variabilidade
milenar
(sub-orbital),
superimposta à primeira, que ocorreu de forma cíclica todos os 1 000 a 1 500
anos (Denton & Karlén, 1973; Koç et al., 1993; O’Brien et al., 1995; Bond et al.,
1997; Campbell, et al., 1998; Bianchi & McCave; 1999; Calvo et al., 2002;
Risebrobakken et al., 2003; Andersen et al., 2004b; Magny, 2004; Mayewski et
al., 2004).
A variabilidade climática holocénica, induzida por modificações dos
parâmetros astronómicos da terra, é geralmente representada por uma subtil
e gradual diminuição da temperatura nos mais variadíssimos registos
sedimentares e de gelo do Hemisfério Norte (Johnsen et al., 2001, Marchal et
al., 2002; Andersen et al., 2004a; Moros et al., 2004). Este padrão gradual de
arrefecimento, foi ainda simulado a partir de modelos numéricos (Kutzbach
& Gallimore, 1988; Crucifix et al., 2002; Weber & Oerlemans, 2003; Renssen et
36
F. Naughton, 2007
al., 2005), nos quais foram introduzidos os valores da insolação de verão a 65
°N definidos para o Holocénico (Berger, 1978).
A data de início deste arrefecimento é, contudo variável, e depende
essencialmente da localização geográfica das zonas estudadas (Kaufman et
al., 2004). De facto, o padrão de arrefecimento gradual tem início durante
ou, após, a ocorrência do episódio térmico máximo do Holocénico (HTMHolocene thermal maximum). Vários registos paleoclimáticos do Atlântico
Norte, detectaram valores de temperatura máxima no início do Holocénico
(11 500 - 8 000 anos cal BP) (Andrews & Giraudeau, 2002; Marchal et al; 2002;
Duplessy et al., 2001; Kaufman et al., 2004; Knudsen et al., 2004; de Vernal et
al., 2005b) enquanto outros, da mesma região e da Gronelândia,
detectaram-no durante ou após 8 000 anos cal BP (no início do Holocénico
médio) (Dahl-Jensen et al., 1998; Bauch et al., 2001; Johnsen et al., 2001;
Levac et al., 2001; Kaplan et al., 2002; Solignac et al., 2004; Kaufman et al.,
2004; Keigwin et al., 2005).
A reconstrução das anomalias da temperatura atmosférica anual
(TANN-annual mean temperature), de inverno (MTCO- mean temperature of
the coldest month) e de verão (MTWA-mean temperature of the warmest
month), efectuada em cerca de 500 sequências polínicas europeias,
permitiram detectar igualmente a ocorrência do HTM, na Europa em geral,
entre 8 000 e 4 000 anos cal BP (Davies et al., 2003). Este autor sugeriu que os
valores da MTWA seriam superiores aos actuais, durante o HTM. O sudoeste
da Europa, incluindo a Península Ibérica, apresenta contudo, valores de
MTWA e MTCO bastante inferiores aos actuais durante o período situado
entre 8 000 e 4 000 anos cal BP (Davies et al., 2003). O limite cronológico
definido por Davies et al. (2003) para o HTM europeu é semelhante ao
intervalo definido a partir de algumas sequências polínicas norte europeias
tais como: nos lagos Raigastvere, Viitna, Ruila (Estónia) e Flarken (Suécia)
(Seppä & Poska, 2004; Seppä et al., 2005).
Para além do gradual decréscimo da temperatura induzido por
modificações dos parâmetros orbitais, a variabilidade climática que ocorreu
à escala milenar está bem expressa em variadíssimos registos climáticos tais
como por exemplo (ver Mayewski et al., 2004):
37
F. Naughton, 2007
-O avanço e recuo de glaciares do norte da Europa e América
(Denton & Karlén, 1973);
- A presença de IRDs e diminuição da SST (de 1 a 3° C) durante os
episódios frios em sondagens marinhas do Atlântico Norte (Bond et al., 1997;
Solignac et al., 2004);
-A subida e descida dos níveis lacustres (ex: Alpes Suissos, Magny et al.,
2003);
-As variações de K+ na sondagem de gelo GISP2, as quais representam
um aumento do fornecimento de poeiras à Gronelândia durante os
episódios mais frios da variabilidade milenar (Mayewski et al., 1997);
-As variações de cor nos sedimentos em sondagens marinhas
recolhidas em “Caríaco Basin” (Hughen et al., 1996).
Dentro desta variabilidade climática milenar, ocorreu um drástico
episódio de arrefecimento, há cerca de 8 200 anos cal BP, o qual é
designado por evento de 8.2 (8.2 kyr event). Este evento foi detectado em
vários registos paleoclimáticos tais como: nas sondagens de gelo da
Gronelândia (O’Brien et al., 1995; Alley et al., 1997; Muscheler et al., 2004),
nas sondagens marinhas do Atlântico Norte (Bond et al., 1997; 2001; Bianchi
& McCave; 1999) e no continente Europeu (Von Grafenstein et al., 1998;
Klitgaard-Kristensen et al., 1998; Nesje & Dahl, 2001; Tinner & Lotter, 2001; 2006;
Baldini et al., 2002; Spurk et al., 2002; Heiri et al., 2003; Magny et al., 2003;
Seppä & Poska, 2004; Veski et al., 2004; Seppä et al., 2005).
Os factores que provocaram este evento têm sido largamente
debatidos nos últimos anos. Grande número de investigadores sugeriu que
este evento teria sido provocado pelo rápido colapso da calote glaciar da
“Laurentide”, que ocorreu há cerca de 8 470 anos (Barber et al., 1999). Este
colapso favoreceu a introdução de grandes quantidades de água doce e
fria no Atlântico Norte, destabilizando o padrão geral da circulação
termohalina e a redução da formação de NADW (Alley et al., 1997; Clark et
al., 2001). A diminuição da intensidade da circulação termohalina impediu o
38
F. Naughton, 2007
transporte de calor para as altas latitudes (Fawcett et al., 1997) provocando
um forte arrefecimento no Atlântico Norte (Barber et al., 1999; Rahmstorf,
2002; Renssen et al., 2001) e, como consequência, terá favorecido a
formação de gelo marinho durante o inverno nas altas latitudes do Atlântico
Norte (Denton et al., 2005). A expansão de gelo marinho amplificou o sinal
de diminuição da temperatura durante o inverno, facilitando assim o
aumento do contraste sazonal anual.
Outros
autores
sugeriram
que
este
evento
foi
provocado
essencialmente por variações relacionadas com a actividade solar (Denton
& Karlén, 1973; Bond et al., 2001; Van Geel et al., 2003). Esta hipótese
pressupõe que as variações da actividade solar produzem flutuações no
vento solar, o qual controla a intensidade dos raios cósmicos e a produção
de
14C
na atmosfera e como tal nos valores de temperatura atmosférica.
Contudo, o facto deste evento ser mais proeminente nos registos climáticos
do Atlântico Norte do que noutras regiões, de ocorrer logo após o colapso
da calote glaciária da “Laurentide”, e do facto de que a reconstrução
baseada em modelos numéricos da anomalia que caracteriza este evento
apresentar
grandes
semelhanças
com
o
sinal
emitido
pelos
vários
indicadores paleoclimáticos (marinhos, continentais e de gelo), vai assim
favorecer a primeira teoria a qual se baseia no designado “freshwater pulse
mechanism” (Alley & Ágústsdóttir, 2005).
Duas publicações recentes, Rohling & Pälike (2005) e Ellison et al.
(2006), sugerem que o evento drástico e abrupto “8.2 kyr event”, ocorreu no
interior de uma longa anomalia fria, de escala secular, entre 8 600 e 8 000
anos cal BP. Este episódio de relativa longa duração (~ 600 anos) foi
observado previamente em alguns trabalhos científicos, nomeadamente:
no registo de K+ da sondagem de gelo GISP2 (Mayeski et al., 1997); nos
resultados de SST obtidos em algumas sondagens marinhas colhidas no
Atlântico Norte (Risebrobakken et al., 2003; Knudsen et al., 2004; Keigwin et
al., 2005) e ainda em sequências polínicas norte europeias (Seppä & Poska,
2004). Contudo este evento foi correlacionado com o evento de 8.2 k anos.
Até à data, poucos estudos foram efectuados de forma a demonstrar
a resposta da vegetação às variações orbitais e sub-orbitais, no sul da
Europa, sendo um exemplo o estudo efectuado por Magri (1995).
39
F. Naughton, 2007
Objectivos 4
Por estas razões, um dos objectivos deste trabalho é testar se a
vegetação das latitudes médias da Europa ocidental respondeu à
variabilidade climática orbital e sub-orbital, e em particular aos eventos de
8.2 k anos e multi-secular (8.6 -8.0 k anos). Pretendeu-se ainda definir o HTM
nesta região e finalmente integrar os resultados e interpretações num
contexto global.
1. 2 Calibração da assinatura polínica marinha ao longo da Península
Ibérica
Tal como foi referido anteriormente, a correlação directa entre os
vários indicadores paleoclimáticos: marinhos, continentais e de volume de
gelo preservados em sedimentos marinhos, permite-nos avaliar o tempo e a
natureza de resposta da vegetação às variações climáticas detectadas no
Atlântico Norte.
Contudo, antes de efectuar este tipo de correlações, deve-se verificar
se os grãos de pólen preservados nas sequências marinhas representam uma
imagem integral da vegetação continental adjacente, numa dada região
em estudo.
Nas últimas décadas, foram elaborados uma série de estudos
palinológicos em amostras sedimentares de superfície, de forma a
determinar os tipos de transporte polínico (do continente para o oceano), os
padrões de dispersão polínica no oceano assim como o tipo de relação
existente entre a assinatura polínica marinha e a sua fonte original, em vários
locais do mundo, nomeadamente: noroeste Pacífico (Heusser & Florer, 1973;
Heusser & Balsam, 1977); no Mar de Okhotsk (noroeste da Russia) (Koreneva,
1966); no Mar do Japão (Koreneva, 1966); no Oceano ĺndico e nos Mares da
Indonésia (van der Kaars, 2001; van der Kaars & de Deckker, 2003); no Mar
mediterrânico (Rossignol, 1969; Koreneva, 1971); no Mar Negro (Koreneva,
1966); no Mar da Noruega (Combaz et al., 1977); a Este do Canadá (Mudie,
40
F. Naughton, 2007
1982); na margem sudoeste francesa (Turon, 1984); no Rio Mississipi (Chmura
et al., 1999); no noroeste da costa Africana (Rossignol-Strick & Duzer, 1979;
Leroy & Dupont, 1994; Hooghiemstra et al., 2006); no Golfo da Guiné (Lézine
& Vergnaud-Grazzini, 1993) e no sudoeste da costa Africana (Dupont &
Wyputta, 2003). A maioria destes estudos mostram que os grãos de pólen
contidos nos sedimentos marinhos dessas regiões representam uma imagem
integral da vegetação regional do continente adjacente.
Sabe-se que os grãos de pólen, após serem produzidos pelas plantas,
são primeiramente dispersos pelo vento, depositados no solo, lagos e rios ou
mantidos a pairar na atmosfera durante um período curto de tempo (Muller,
1959; Cour et al., 1999). O meio de transporte polínico desde as zonas
costeiras para o mar aberto depende essencialmente das condições
ambientais de cada zona (Groot & Groot, 1966; Koreneva, 1966; Dupont et
al., 2000):
- em zonas áridas, com sistemas hidrológicos pouco desenvolvidos, tais
como o noroeste da costa Africana (Leroy & Dupont, 1994; Hooghiemstra et
al., 2006) e o Canadá oriental (Mudie & McCarthy, 1994), o transporte
polínico do continente para o oceano, é essencialmente efectuado pelo
vento;
- em zonas áridas a semi-áridas, atravessadas por alguns sistemas
fluviais tais como o Golfo da Guiné (Lézine & Vergnaud-Grazzini, 1993) e o
Mar Alboran (Moreno et al., 2002), o tipo de transporte polínico é misto
(fluvial e eólico);
- em zonas costeiras com sistemas fluviais complexos e bacias
hidrográficas bem desenvolvidas, os grãos de pólen são essencialmente
transportados para o mar pelos rios (Muller, 1959; Cross et al., 1966; Bottema
& Van Straaten, 1966; Heusser & Balsam, 1977; Peck, 1973).
Vários trabalhos experimentais, efectuados desde a década de 60,
mostraram que os grãos de pólen e esporos são considerados como
partículas sedimentares finas, que quando suspensos na água comportam-se
41
F. Naughton, 2007
e obedecem às mesmas leis físicas dos sedimentos finos em suspensão
(Muller, 1959; Koreneva, 1966; Traverse & Ginsburg, 1966; Stanley, 1966; Groot
& Groot, 1966; Chmura & Eisma, 1995). Os grãos de pólen, tal como as
partículas
sedimentares
finas,
são
ainda
sujeitos
a
processos
de
aglomeração, floculação e pelitização, que favorecem o aumento da
velocidade de sedimentação dos mesmos em meios marinhos (Chmura &
Eisma, 1995; Chmura et al., 1999). Vários autores estimaram que a velocidade
de sedimentação na coluna de água é cerca de 100 m/dia (Hogghiemstra
et al., 1992).
Posteriormente a serem libertados no mar pelos rios ou pelo vento, os
grãos de pólen podem ser depositados na plataforma continental ou ser
transportados através de correntes oceânicas para o mar profundo (Muller,
1959; Heusser & Shackleton, 1979; Muller, 1959; Groot & Groot, 1966; Dupont
et al., 2000). Estes podem ser transportados para distâncias longínquas à
linha de costa tais como 1500 a 2500 Km (Muller, 1959). Contudo, alguns
autores sugerem que, distâncias à linha de costa superiores a 500 Km
apresentam geralmente um fraco conteúdo esporopolínico, e que o
espectro polínico dessas sequências pode, por vezes, não representar uma
imagem integral da vegetação do continente adjacente (Koreneva, 1966;
Stanley, 1966).
Em meio marinho, o conteúdo polínico diminui gradualmente à
medida que nos distanciamos da actual linha de costa (em direcção ao
largo), seguindo geralmente o mesmo padrão de dispersão das partículas
sedimentares finas (Muller, 1959; Cross et al., 1966; Groot & Groot, 1966;
Bottema & Van Straaten, 1966; Koreneva, 1966; Stanley, 1966; van der Kaars
& de Deckker, 2003).
Durante a última década, vários trabalhos de palinologia foram
realizados em sondagens marinhas profundas ao longo da margem Ibérica
de forma a compreender a relação entre as modificações do coberto
vegetal a as variações climáticas detectadas no Atlântico Norte e
Gronelândia (Hooghiemstra et al., 1992; Sánchez Goñi et al., 1999; 2000; 2002;
2005; Boessenkool et al., 2001; Roucoux et al., 2001; 2005; Turon et al., 2003;
Tzedakis et al., 2004; Desprat, 2005; Desprat et al., 2005; 2006; in press).
42
F. Naughton, 2007
Contudo, e até à data, não foram efectuados quaisquer tipo de
estudos experimentais de forma a demonstrar que os grãos de pólen
preservados nessas sequências marinhas representam uma imagem integral
da vegetação da Península Ibérica, e estudos que permitam determinar os
tipos de mecanismos envolvidos na transferência desses grãos, do
continente para o oceano, nesta região.
Objectivo 5 Pretende-se desta forma:
- investigar se a assinatura polínica das amostras de superfície costeiras e
marinhas do noroeste e sudoeste da Península Ibérica é semelhante aos
espectros polínicos actuais representantes das regiões biogeográficas
Atlântica e Mediterrânica, respectivamente;
- mostrar que a assinatura polínica marinha actual representa uma imagem
integral regional e não uma imagem local da vegetação do continente
adjacente;
- determinar os padrões de transporte e dispersão polínica, do continente
para o mar aberto, a partir do estudo comparativo entre a concentração
polínica total de cada amostra e modelos de dispersão sedimentar de
partículas finas efectuados ao longo da margem ibérica.
43
F. Naughton, 2007
1. 3 Impacto da variabilidade climática na evolução dos sistemas
costeiros
As variações climáticas têm um papel muito importante no controlo
das modificações do nível do mar global, as quais afectam profundamente
os sistemas costeiros.
1. 3. 1 Variações do nível médio do mar
Tal como foi referido anteriormente, durante o Quaternário, o clima
terrestre sofreu variações de grande e pequena escala induzidas por
modificações internas e ou externas (insolação e a constante solar) que
ocorreram no planeta. Estas variações estão associadas a mudanças mais
ou menos importantes no volume de gelo acumulado nos pólos e
continentes que, por consequência, provocaram alterações no nível médio
do mar (Shackleton & Opdyke, 1973).
As modificações do nível do mar são, contudo, espacialmente
afectadas por oscilações gravitacionais potenciais do sistema terra-oceano
e gelo, sendo por isso necessário efectuar correcções a nível glacio-hidroisostático durante a elaboração de uma curva de variação do nível do mar
global (Lambeck et al., 2002).
Nos últimos anos, várias curvas de variação do nível do mar foram
elaboradas na zona tropical, tendo em conta as variações de altura dos
terraços coralíferos datados a partir do U/Th obtidos em: Barbados
(Fairbanks, 1989; Bard et al., 1990a; 1990b), Tahiti (Bard et al., 1996), Península
de Huon (Chappell & Polach, 1991; Edwards et al., 1993), noroeste
Australiano (Yokoyama et al., 2000; 2001) etc. e utilizadas por Lambeck et al.
(2002) para estimar a variabilidade do nível do mar global e respectivas
modificações no volume de gelo desde o estádio isotópico marinho 3 (MIS 3)
até à actualidade (Fig. I.14).
O nível do mar desceu desde 60 000 a 30 000 anos cal BP onde atingiu
o seu valor mínimo de – 120 m. Entre 30 000 e 19 500 anos, o nível do mar
permaneceu relativamente estável e, no final deste período, subiu
rapidamente (15 m) em cerca de 500 anos (Yokoyama et al., 2000).
44
F. Naughton, 2007
Fig. I.14 | Curva de variação do nível global do mar (adaptado de Lambeck et al., 2002).
Durante o início da deglaciação (entre 19 000 e 16 000 anos) a fusão
global dos gelos foi lenta provocando um aumento gradual do nível do mar
de cerca de 3.3 mm/ano, enquanto que entre 16 000 e 12 500 anos cal BP a
subida do nível do mar foi bastante mais rápida (16.7 mm/ano) (Lambeck et
al., 2002). Há cerca de 14 000 anos, um hiato, marca o evento designado de
“melwater pulse 1A” (Fairbanks, 1989; Bard et al., 1990a). A ocorrência de um
episódio de curta duração, entre 12 500 e 11 500 anos cal BP, marcado por
um “plateau” onde o nível do mar permaneceu relativamente estável,
representa o evento Dryas recente (Bard et al., 1996; Lambeck et al., 2002).
Desde 11 500 a 8 500 anos o nível do mar subiu gradualmente cerca de 15.2
mm/ano (Lambeck et al., 2002). Finalmente o volume oceânico aproximouse do seu actual valor há cerca de aproximadamente 7 000 anos (Lambeck,
2000; Lambeck et al., 2002).
1. 3. 2 Evolução dos sistemas costeiros
As modificações do nível do mar que ocorreram durante a
deglaciação, encontram-se registadas nas sequências sedimentares por
uma alternância entre episódios de incisão fluvial ou de preenchimento
sedimentar, associados a momentos de descida ou subida do nível do mar,
respectivamente. Estas características foram detectadas pela primeira vez
45
F. Naughton, 2007
no Vale do Mississipi (Fisk & McFarlan, 1955) e posteriormente noutros sistemas
costeiros mundiais, nomeadamente: na costa da Louisiana (Nichol et al.,
1996), na costa de Delaware (Belknap & Kraft, 1985), no sul da Austrália (Roy
et al., 1995) e no delta Yangtze na China (Li et al., 2002).
Para além deste indicador sedimentar de modificações relativas do
nível relativo do mar, existem também outro tipo de variações, relacionadas
com a amplitude da influência fluvial relativamente à marinha e vice-versa,
nos registos sedimentares costeiros, os quais podem fornecer informações
importantes no que respeita à variabilidade do nível de base marinho
(Dalrymple et al., 1994; Zaitlin et al., 1994).
Este tipo de abordagem, que se baseia no estudo das mudanças
ambientais ao longo de registos sedimentares costeiros (lagoas e estuários)
durante a deglaciação, foi efectuado nos últimos anos ao longo da
Península Ibérica nomeadamente na zona costeira sudoeste portuguesa
(Freitas et al., 2002; 2003). Estes trabalhos sugerem a presença de um
ambiente fluvial durante o LGIT seguido de um gradual aumento da
influência marinha ao longo do Holocénico como resposta à subida gradual
do nível do mar o qual atingiu a sua máxima influência no final da
deglaciação. Contudo, no final da deglaciação, a formação de barreiras
arenosas na embocadura destes estuários assim como a desaceleração da
subida do nível do mar contribuíram, para que outros mecanismos locais
tenham sido responsáveis pelas modificações geomorfológicas destas zonas
(Freitas & Andrade, 1997; 2001; Freitas et al., 2002; 2003).
Este tipo de estudos começou a ser efectuado para a região noroeste
de Portugal nos últimos anos (Naughton, 2002; Guerreiro et al., 2005; Moreno
et al., 2005; 2006; Drago et al., in press; Fradique et et al., in press).
De forma a compreender a evolução geomorfológica do noroeste de
Portugal e descriminar os mecanismos globais e locais responsáveis pela
evolução do estuário do Douro, foi efectuado um estudo polínico e
sedimentar, de alta resolução temporal, em duas sequências sedimentares
estuarinas situadas no noroeste da Península Ibérica para o Holocénico
(Naughton, 2002). Este trabalho foi sintetizado numa publicação científica, a
qual é apresentada no capítulo 6.
46
F. Naughton, 2007
1. 4 Zona de estudo
1. 4. 1 Margem Ibérica
1. 4. 1. 1 Clima e vegetação actual
A Península Ibérica é ocupada por uma série de biomas, climas e tipos
de solos que variam de acordo com a topografia. Esta região é
caracterizada por duas zonas biogeográficas principais: a zona atlântica
(Blanco Castro et al., 1997) e a zona mediterrânica (Fig. I.15). A zona
atlântica inclui, a norte da Península ibérica, a designada zona Eurosiberiana
(Peinado & Rivas-Martinez, 1987) (Fig. I.10).
Uma vez que os pólens e esporos inclusos nos sedimentos marinhos
representam geralmente a vegetação que coloniza as bacias hidrográficas,
deve-se ter em conta ambas as delimitações biogeográficas referenciadas
nas figuras seguintes: Fig. I.10 e Fig. I.15.
Fig. I.15 | Localização das zonas biogeográficas (adaptado de Blanco Castro et al., 1997).
A região biogeográfica mediterrânica ocupa grande parte da
Península enquanto que a zona atlântica é estreita e está essencialmente
confinada ao norte e noroeste da mesma. A região mediterrânica é
caracterizada por um forte contraste sazonal (verões quentes e secos e
47
F. Naughton, 2007
invernos frios e húmidos) enquanto a região atlântica, influenciada pelo
oceano Atlântico, apresenta condições sazonais menos contrastantes,
sendo por isso caracterizada por um clima temperado e húmido ao longo de
todo o ano.
O noroeste espanhol, incluindo a bacia hidrográfica do Minho, é
caracterizado por temperaturas situadas entre -7 e +10°C e por uma
precipitação média anual que varia entre 900 e 1400 mm. Esta zona é
dominada pela floresta de carvalhos, nomeadamente por: Quercus robur,
Q. pyrenaica e Q. petraea e por uma vegetação rasteira constituída
essencialmente por herbáceas do tipo Ericaceae, Calluna e Ulex. Nesta
zona, é ainda possível detectar localmente Betula pubescens subsp.
celtiberica, Corylus avellana e Genista (Alcara Ariza et al., 1987).
Ligeiramente a sul desta região, encontra-se uma zona de transição a
qual inclui a bacia hidrográfica do Douro. Esta zona é caracterizada por
valores de temperatura média anual de 12º C (Loureiro et al., 1986) e por
temperaturas invernais que rondam os 4 e -4°C nas baixas e médias altitudes,
podendo chegar aos -8°C nas altas altitudes (Polunin & Walters, 1985). A
precipitação é elevada, sendo aproximadamente de 700 a 1000 mm/ano
nas baixas e médias altitudes, e de 1600 mm/ano nas altas altitudes. A
influência oceânica é particularmente importante no noroeste da bacia,
favorecendo a predominância da associação Q. robur e Q. suber (BraunBlanquet et al., 1956). A expansão de Pinus pinaster, Pinus sylvestris,
Eucalyptus globulus, Castanea sativa e Juglans regia resulta do impacto
antrópico (Valdès & Gil Sanchez, 2001). A vegetação é ainda dominada por
Ulex e Ericaceae e as margens do rio são colonizadas por Alnus glutinosa,
Fraxinus angustifolia, Ulmus spp., Salix spp. e Populus spp.. A zona oriental é
afectada por condições ligeiramente mais continentais, favorecendo a
expansão de Q. ilex, Q. suber, Juniperus spp. e a floresta de carvalhos (Q.
pyrenaica, Q. faginea).
O sudoeste da Península Ibérica, nomeadamente as bacias do Tejo e
Sado, são influenciadas por um clima mediterrânico (precipitação média
anual de 200 a 600 mm e temperaturas que variam entre 4 a 14°C) o qual
favorece o desenvolvimento da floresta esclerófila. A zona oeste é
representada essencialmente por uma floresta de Q. Ilex, Q. rotundifolia e Q.
48
F. Naughton, 2007
Suber, assim como pela presença de Phillyrea angustifolia e Pistacia
terebinthus enquanto a parte oriental, é sobretudo colonizada por Q.
rotundifolia e Q. coccifera associada a Juniperus communis e Pinus
halepensis. As zonas representantes de altitudes médias (700-1000 m a.s.l.)
são dominadas pela floresta de carvalhos de folha caduca (Q. pyrenaica e
Q. faginea) associada a espécies norte Europeias tais como Taxus baccata.
A degradação da floresta favorece a expansão de Cistaceaes nas zonas
ligeiramente áridas e de Ericaceas nas zonas mais húmidas (Blanco Castro et
al., 1997).
1. 4. 1. 2 Oceanografia
A margem Ibérica é dominada pelo sistema de correntes superficiais
designado por Sistema de correntes de Portugal (PCS-Portugal Current
System) o qual é composto por uma corrente lenta que se desloca em
direcção ao equador ao longo do oceano aberto (Arhan et al., 1994) (Fig.
I.16) e pela rápida corrente costeira a qual reverte sazonalmente a direcção
do seu percurso (Ambar & Fiúza, 1994; Barton, 1998).
Fig. I.16 | Esquema detalhado das principais correntes de superfície do Atlântico Norte: EG-corrente Este
da Gronelândia, Ei-corrente este da Islândia, Gu-Gulf Stream, Ir-corrente de Irminger, La-corrente do
Labrador, Na-corrente Norte Atlântica, Nc-corrente do Cap Norte, Ng-corrente da Noroega, Ni-corrente
do Norte da Islândia, Po-corrente de Portugal, Sb-corrente de Spitsbergen, Wg-corrente Oeste da
Gronelândia. Linhas negras representam as correntes relativamente quentes enquanto as linhas a
tracejado representam correntes relativamente frias (adaptado de Dietrich et al., 1980).
49
F. Naughton, 2007
Durante o verão, a célula de altas pressões dos Açores encontra-se
localizada na zona central do Atlântico Norte e o centro de baixas pressões
da Gronelândia é fraco (Torres et al., 2003) (Fig. I.17).
Fig. I.17 | Representação da localização dos centros de altas e baixas pressões durante o verão e o
Inverno ao longo do Hemisfério Norte (adaptado de Hurrell & Dickson, 2004). As setas representam a
direcção dos ventos dominantes.
Esta situação gera ventos dominantes vindos de norte e do noroeste
favorecendo a circulação para sul das correntes costeiras superficiais (Fiúza
et al., 1982; Haynes & Barton, 1990), ao longo da zona externa da
plataforma, nos primeiros 50 a 100 m de profundidade da coluna de água
(Álvarez-Salgado et al., 2003), assim como a ocorrência de upwelling ao
longo da margem portuguesa (Fiúza et al., 1982; Haynes & Barton, 1990;
Torres et al., 2003). A massa de água fria rica em nutrientes (ENACWspEastern North Atlantic Central Water of subpolar sources) dirige-se para norte
dos 45° N enquanto a massa de água salina e pobre em nutrientes
(ENACWst- Eastern North Atlantic Central Water of subtropical origin) dirigese0 para sul de 40° N (Fiúza, 1984; Rios et al., 1992) (Fig. I.18).
Durante o inverno, o centro de altas pressões dos Açores localiza-se
na margem noroeste Africana, o centro de baixas pressões da Gronelândia
é forte e encontra-se desviado para sudoeste da Gronelândia (Fig. I.17).
50
F. Naughton, 2007
Fig. I.18 | Esquema das principais correntes oceânicas que circulam ao longo da margem Ibérica. PCS:
Portugal Current system; ENACWsp: Eastern North Atlantic Central Water de origem sub-polar; ENACWst:
Eastern North Atlantic Central Water de origem sub-tropical; MSW: Mediterranean Sea Water; LSW: Labrador
Sea water; NADW: North Atlantic Deep Water (adaptado de Sprangers et al., 2004).
O gradiente de pressão entre estes dois sistemas depressionários
favorece os ventos vindos de sul, ao longo da margem ibérica, provocando
processos de downwelling e o transporte das correntes costeiras de
superfície para norte (Frouin et al., 1990; Haynes & Barton, 1990). Esta inversão
dos padrões hidrológicos ocorre entre Setembro-Outubro e Março-Abril e
representa a designada Contra-corrente costeira portuguesa (PCCCPortugal Coastal Counter Current) (Ambar et al., 1986). Esta circulação
ocorre numa zona estreita (aproximadamente de 30 Km) e transporta uma
massa de água quente e salina (ENACWst) para norte, a profundidades
compreendidas entre 200 a 300 m (Pingree & Le Cann, 1990).
Entre 550 a 1500 m de profundidade da coluna de água, a massa de
água altamente salina e relativamente quente (MSW- Mediterranean Sea
Water) proveniente do estreito de Gibraltar migra para norte (Mazé et al.,
1997) (Fig. I.18). Contudo, a salinidade desta massa de água diminui
principalmente a latitudes superiores a 41° N onde se mistura com a massa
de água de fraca salinidade proveniente do Mar de Labrador (LSWLabrador Sea water) (McCave & Hall, 2002). A LSW é uma das três
componentes que compõem a massa de água profunda norte Atlântica
(NADW- North Atlantic Deep Water) (Huthnance et al., 2002).
51
F. Naughton, 2007
1. 4. 1. 3 Geomorfologia e dinâmica sedimentar actual
A Margem Ibérica é caracterizada por uma plataforma continental
muito estreita (30 a 50 Km) e por um talude irregular, bastante inclinado, o
qual mergulha rapidamente para as grandes profundidades da planície
abissal (Fig. I.19). Esta margem continental é interceptada parcialmente por
canhões submarinos muito profundos, nomeadamente: Mugia, Porto, Aveiro,
Cascais, Lisboa e São Vicente. Os grandes canhões submarinos de Setúbal e
da Nazaré cortam a plataforma continental na sua quase totalidade
facilitando a captura dos sedimentos que se encontram em suspensão
permitindo assim a sua conduta directa para a planície abissal (Vanney &
Mougenot, 1981). Pensa-se que os canhões submarinos Ibéricos devem ter
sido bastante mais activos durante períodos de baixo nível do mar (Van
Weering & McCave, 2002). O talude continental é ainda marcado pela
presença de várias montanhas submarinas tais como: Vigo (VS), Vasco da
Gama (VDGS), Porto (PS), Tore (TS), pelo Banco da Galiza e por algumas
depressões tectónicas (Vanney & Mougenot, 1981).
Fig. I.19 | a) Morfologia da margem continental Ibérica: Canhões submarinos de Mugia (MC), do Porto
(PC), de Aveiro (AC), da Nazaré (NC), de Cascais (CC), de Lisboa (LC), de Setúbal (SC) e de São Vicente
(S.VC); montanhas submarinas de Tore (TS), do Porto (PS), Vasco da Gama (VDGS) e de Vigo (VS).
52
F. Naughton, 2007
No noroeste da Península Ibérica, grandes quantidades de sedimentos
são libertados pelos rios Douro, Ave, Cavado, Lima e Minho para a margem
continental Ibérica. O principal fornecedor de sedimentos à plataforma
noroeste portuguesa é o rio Douro seguido pelo rio Minho (~8.2 x 109m3
descarga média anual) (Dias et al., 2002; Jouanneau et al., 2002; Oliveira et
al., 2002) (Fig. I.20). A área das suas bacias hidrográficas é de 97 682 km2 e 17
081 km2 e o seu comprimento total é de 927 km e 300 km, respectivamente
(Loureiro et al., 1986). Contrariamente aos rios que desaguam no noroeste de
Portugal, as rias situadas a norte de 42° N (Vigo, Pontevedra, Arousa e Muros)
são consideradas como armadilhas sedimentares impedindo o fornecimento
de sedimentos continentais à plataforma adjacente (Dias et al., 2002;
Jouanneau et al., 2002).
A plataforma continental noroeste Ibérica é composta por: a) uma
zona interna, situada a menos de 30 m de profundidade, constituída
essencialmente por areias finas; b) uma zona intermédia constituída
principalmente por areias e cascalhos e uma zona externa rica em
carbonatos e areias de dimensão média (Van Weering et al., 2002). Esta
plataforma é ainda constituída por dois importantes complexos silto-argilosos
(Douro e Galiza) separados por uma zona desprovida de lodo (Lopez-Jamar
et al., 1992) (Fig. I.20).
Fig. I.20 | Morfologia da plataforma continental do noroeste da Península Ibérica (adaptado de Dias et al.,
2002).
53
F. Naughton, 2007
A evolução destes corpos lodosos depende essencialmente da
quantidade de sedimentos fornecidos pelo continente adjacente, da acção
das barreiras morfológicas e ainda, das condições hidrológicas do meio
(Dias et al., 2002; Jouanneau et al., 2002).
Os processos de sedimentação são bastante complexos no noroeste
da margem Ibérica e ocorrem muitas vezes associados a eventos episódicos
de cheias (Dias et al., 2002) e /ou a episódios associados a descargas fluviais
importantes (Araújo et al., 1994; Drago et al., 1998).
Depois de serem libertados pelos rios, os sedimentos finos são
transportados essencialmente por camadas nefelóides: de fundo (BNLBottom Nepheloid Layers), intermédias (INL-Intermediate Nepheloid Layers) e
de superfície (SNL-Surface Nepheloid Layers). Oliveira et al. (1999),
detectaram um decréscimo da concentração de sedimentos finos ao longo
das diferentes camadas nefelóides com o aumento da distância à linha de
costa e que as correntes e a ondulação induzem a resuspensão dos
sedimentos depositados nos complexos silto-argilosos do Douro e Galiza
especialmente durante episódios de grandes tempestades. Durante alguns
destes episódios extremos, tais como tempestades de inverno as quais estão
associadas a condições de downwelling (ventos vindos do sul), os
sedimentos contidos na BNL podem ficar retidos nos afloramentos rochosos
(Fig. I.20) (Drago et al., 1998) os quais funcionam como barreiras à
transferência de sedimentos da plataforma para o talude e planície abissal
(Jouanneau et al., 2002; Van Weering et al., 2002), estimulando assim o
transporte sedimentar para norte, ao longo da plataforma continental
(Drago et al., 1998; Dias et al., 2002; Jouanneau et al., 2002; Van Weering et
al., 2002).
A
ondulação
associada
a
esses
eventos
extremos
podem
esporadicamente induzir a ressuspensão de partículas situadas na BNL
(Vitorino et al., 2002) alimentado a camada subjacente INL (Oliveira et al.
2002) e contribuir para a exportação dos mesmos para o largo (Vitorino et
al., 2002).
A inversão de correntes de superfície induzidas pela presença local de
redemoinhos pode também contribuir para a transferência de partículas
para o talude e planície abissal (Pingree & LeCann, 1992).
54
F. Naughton, 2007
Durante o verão e associado a condições de upwelling (ventos vindos
do norte ou de noroeste) a exportação de sedimentos efectua-se
essencialmente a nível do eixo da plataforma (McCave & Hall 2002; Van
Weering et al., 2002). A transferência lateral de sedimentos da plataforma
para o largo pode ocorrer por vezes, em locais onde existam filamentos
transversais (Huthnance et al., 2002).
A sudoeste da margem Ibérica, o rio Tejo é o principal fornecedor de
sedimentos à plataforma continental, talude e planície abissal, seguido pelo
rio Sado (Dias, 1987; Jouanneau et al., 1998) (Fig. I.21). O Rio Tejo apresenta
cerca de 1110 km de comprimento e a área da sua bacia hidrográfica é de
80 600 km2 enquanto o rio Sado é bastante mais pequeno, com apenas 175
km de comprimento e uma área da bacia hidrográfica de 7 640 km2
(Loureiro et al., 1986).
Fig. I.21 | Morfologia da plataforma continental sudoeste portuguesa (adaptado de Araújo et al., 2002).
A diferença entre a descarga fluvial existente entre os dois rios assim
como o papel desempenhado pelas correntes litorais, condicionam a
dinâmica sedimentar desta zona (Jouanneau et al., 1998). O complexo silto-
55
F. Naughton, 2007
argiloso localiza-se logo após a embocadura do estuário do Tejo e cobre
toda a área da plataforma continental (Araújo et al., 2002).
Durante o verão, a concentração de matéria particulada em
suspensão (SPM-Suspended Particulate Matter) é quatro vezes mais elevada
na embocadura do estuário do Tejo do que no Sado e as camadas
nefelóides estendem-se cerca de 30 Km para oeste da actual linha de costa
(Jouanneau et al., 1998).
O transporte de partículas finas para o largo é favorecido pelos
canhões submarinos de Cascais, Lisboa e Setúbal (Jouanneau et al., 1998) e
por filamentos transversais (Huthnance et al., 2002).
56
F. Naughton, 2007
1. 4. 2 Plataforma continental noroeste Francesa
1. 4. 2. 1 Geomorfologia, oceanografia e sedimentação actual
A plataforma continental Francesa localiza-se no Golfo da Gasconha
(Golfe de Gascogne) entre 43° a 48° N (Fig. I.22). Morfologicamente a
plataforma francesa é muito larga e pouco inclinada no noroeste (300 Km
de largura) tornando-se dez vezes mais estreita e íngreme para sul (30 Km de
largura).
Fig. I.22 | Localização do corpo lodoso “Grande Vasière” na plataforma continental Francesa.
Esta plataforma, é constituída por dois corpos lodosos principais: um
situado próximo do estuário da Gironde (Gironde shelf mud field) e um outro,
situado entre a Bretanha e a Charente (Grande Vasière) (Allen & Castaing,
1977). Estes corpos lodosos são essencialmente alimentados por sedimentos
finos libertados pelos rios Gironde e Loire e, em menores quantidades, pelos
rios Adour, Vilaine e Charente (Castaing & Jouanneau, 1987; Jouanneau et
al. 1999). A área das suas bacias hidrográficas é de: Gironde (que inclui as
bacias hidrográficas dos rios Garonne e Dordogne) 79 180 km2, Loire 118 420
km2, Adour 17 100 km2, Vilaine 10 420 km2 e Charente 11 970 km2.
O estuário da Gironde é um dos mais longos da Europa, apresentando
cerca de 80 km de comprimento e uma superfície que ronda os 625 Km2.
A plataforma Francesa é afectada fortemente pela ondulação,
principalmente em períodos de fortes tempestades. Para além da
57
F. Naughton, 2007
ondulação, esta plataforma é ainda sujeita a um regime de maré semidiurno, meso a macrotidal. Durante o verão a massa de água apresenta
uma forte estratificação enquanto que, durante o inverno, a acção dos
ventos que sopram de Oeste favorecem a não estratificação da coluna de
água impedindo a exportação de sedimentos finos para o talude e planície
abissal.
O complexo silto-argiloso “Grande Vasière”, apresenta cerca de 250
Km de comprimento (paralelamente à linha de costa actual), situa-se a
cerca de 100 m de profundidade da coluna de água, e apresenta uma taxa
de sedimentação de 0.1-0.2 cm/ano (Lesueur et al., 2001) (Fig. I.22).
O fornecimento de sedimentos através das camadas nefelóides
permite a manutenção deste corpo lodoso (Jouanneau et al:, 1999; Lesueur
et al., 2001), principalmente durante períodos de cheia (Bourillet et al., 2005).
Durante estes episódios cerca de 30% dos sedimentos em suspensão são
libertados pelos rios e depositam-se a cerca de 20 Km da costa.
1. 4. 2. 2 Clima e vegetação
O noroeste francês é caracterizado por um clima teperado e húmido,
com temperaturas médias anuais que rondam os 13°C e precipitação média
anual de cerca de 1000 mm (dados obtidos na agência pública: “Meteo
France”).
À medida que nos afastamos do litoral, a precipitação média anual
diminui até cerca 600 mm/ano, nas zonas de baixa altitude, enquanto que
estes valores são bastante elevados em áreas situadas nas altas altitudes, tais
como: o “Massif Central”, que apresenta valores superiores a 2200 mm/ano e
nos Pirinéus onde variam entre 2000 mm/ano (oeste) e 1000 mm/ano (este).
O clima temperado e húmido que caracteriza o noroeste francês,
favorece a expansão da floresta decídua e da floresta mista quente,
nomeadamente: deciduous Quercus (Quercus pedunculata, Q. pubescens e
Q. sessiflora) e a presença local de Q. Ilex, Q. Suber, Ulmus e Fraxinus. As
zonas litorais são essencialmente compostas por Pinus pinaster e Ulex. Nas
zonas de maiores altitudes é possível encontrar ainda Fagus e Carpinus.
58
F. Naughton, 2007
1. 5 Material e Métodos
1. 5. 1 Amostras superficiais
De forma a calibrar a assinatura polínica marinha ao longo da
Península Ibérica, foram recolhidos uma série de níveis sedimentares
superficiais, em vários tipos de ambientes distintos, a oeste da Península
Ibérica, tais como: estuários, plataforma e talude continental (Fig. I.23; Tab.
I.1), nos quais foi efectuada a análise do conteúdo polínico. Algumas destas
amostras foram recolhidas na área adjacente à zona biogeográfica
mediterrânica, enquanto outras foram preferencialmente recolhidas na
região
adjacente
à
zona
biogeográfica
Atlântica.
Cada
amostra
corresponde a um sinal polínico representativo dos últimos 200 anos.
Fig. I.23 | Amostras sedimentares de superfície analisadas neste estudo (VIR-18, Ría de Vigo; Laquasup,
Estuário do Douro; PO287-13-2G, Complexo silto-argiloso do Douro; CG11, Complexo silto-argiloso do
Minho; MD99-2331, Talude continental ao largo de Vigo; MD04-2814 CQ, Talude continental ao largo do
Porto; Barreiro, Estuário do Tejo; MD99-2332, Complexo silto-argiloso de Lisboa; FP8-1, Talude continental
ao largo de Sines; e MD95-2042, Talude-planície abissal ao largo de Sines).
59
F. Naughton, 2007
Nome da
amostra
Prof.
(cm)
Latitude
Longitude
MD95-2042
Top
37°48’N
Ano
10°10’W
Prof.
coluna
água
(m)
3148
1995 IMAGES V
38°01’N
09°20’W
980
2003 FORAMPROX1
38°33’N
09°22’W
97
1999 GINNA-IMAGESV
38°40’N
09°07’W
0
1999 IPIMAR
40°37’N
09°52’W
2449
2004 ALIENOR
41°09’N
08°38’W
0
41°09’N
09°01’W
81
41°48’N
09°04’W
107
2001 LAQUATEDE
(PRAXIS/P/CTE/1
1101/1998)
ENVI-CHANGES
(PDCTM/PP/MAR
/15251/99)
2002 ENVI-CHANGES
PDCTM/PP/MAR/
15251/99
1992 GEOMAR92
Projecto,
missões
ou instituições
(0-1)
1FP8-1
Top
(0-1)
MD99-2332
Top
(0-1)
Barreiro
Top
(0-1)
MD04-2814 CQ Top
(0-1)
Laquasup
Top
(0-5)
PO287-13-2G
Top
(0-1)
CG11
Top
(0-1)
MD99-2331
3-4
42°09’N
09°42’W
2120
Vir-18
Top
42°14’N
08°47’W
45
(0-1)
1999 GINNA-IMAGES
V
1990 Departamento
de estratigrafia
da Universidade
de Vigo
Tab. I.1 | Localização das amostras de superfície. Da esquerda para a direita encontra-se representado o
nome das amostras, a profundidade na coluna sedimentar, a latitude, a longitude, a profundidade da
coluna de água, o ano da colheita das amostras, o nome dos projectos científicos ou o nome das missões
oceanográficas ou o nome das Instituições que forneceram as amostras.
Os resultados obtidos foram posteriormente comparados com as
assinaturas
polínicas
de
várias
amostras
de
superfície
continentais
(sedimentos lacustres, musgos e turfeiras) provenientes da base de dados
“European
pollen
database”
http:/www.imep-cnrs.com/pages/EPD.htm
(Peyron et al., 1998; Barboni, et al., 2004). Estas amostras representam
igualmente um sinal polínico dos últimos 200 anos.
Na tentativa de compreender as variações do registo polínico obtido
em duas sequências sedimentares colhidas no estuário do Douro (Core 1 e
60
F. Naughton, 2007
Core 1B) (Capítulo 6), foi necessário, efectuar um estudo polínico em cinco
amostras de superfície colhidas após um episódio de cheia que ocorreu em
Novembro de 2000, de forma a poder descriminar os diferentes tipos de
assinatura polínica: regional (associado à cheia) e local (associada a uma
parcial exposição sub-aérea) (Fig. I.24).
Fig. I.24 | Localização das amostras de superfície e das sondagens estudadas ao longo do estuário do
Douro. Os círculos brancos com pinta negra representam as amostras de superfície e os círculos pretos
representam as sondagens.
61
F. Naughton, 2007
1. 5. 2 Sondagens
Duas sondagens marinhas profundas (MD99-2331 e MD03-2697), uma
sondagem marinha pouco profunda (Vk03-58Bis) e duas sondagens
estuarinas foram escolhidas para estudar a variabilidade climática dos
últimos 30 000 anos nas latitudes médias do Atlântico Norte (Fig. I.25).
Fig. I.25 | Localização das sondagens utilizadas neste trabalho. Sondagens marinhas profundas: MD992331 e MD03-2697; sondagem marinha pouco profunda: VK03-58Bis; sondagens estuarinas: Core1 e
Core1B.
1. 5. 2. 1 Sondagens marinhas profundas
As sondagens MD99-2331 (42° 09’ N, 09° 40’ 90 W) e MD03-2697 (42° 09’
59 N, 09° 42’ 10 W) foram colhidas na margem Galega, a 2110 m e 2164 m de
profundidade, a bordo do navio oceanográfico Marion Dufresne, usando um
“corer” CALYPSO, durante as missões oceanográficas GINNA (a qual se
insere no programa IMAGES V) e PICABIA, respectivamente (Fig. I.25). Estas
sondagens apresentam cerca de 37.2 m e 41.23 m sedimento hemipelágico
e cobrem aproximadamente os últimos 225 000 e 425 000 anos.
195 níveis sedimentares foram recolhidos em todos os 2 cm nos
primeiros 3 m da sondagem MD99-2331 e, em todos os 5 cm nos restantes 4
m de sedimento. Infelizmente, os resultados obtidos ao longo dos níveis
sedimentares correspondentes ao MIS 1 (primeiros 2 metros de sondagem), o
qual engloba o LGIT e o actual interglaciário (Holocénico), não serão
62
F. Naughton, 2007
apresentadas nesta dissertação uma vez que, posteriormente às análises
efectuadas (pólen, δ18O de foraminíferos planctónicos e bentónicos,
alcanonas, associações de foraminíferos planctónicos), verificou-se a
existência de mistura de sedimentos, nos resultados obtidos pelos vários
indicadores paleoclimáticos, assim como, nos resultados obtidos pelas
idades 14C. Esta mistura de sedimentos, foi igualmente confirmada a partir da
análise radiográfica a qual foi efectuada utilizando um processador de
Imagem (SCOPIX image-processing mode). Por essa razão, apenas 110 dos
níveis analisados (a nível do seu conteúdo polínico) serão apresentados
neste trabalho.
De forma a complementar os resultados obtidos na sondagem MD992331 foi efectuada a análise polínica em 54 amostras, ao longo dos primeiros
4.10 m de sedimento, na sondagem vizinha MD03-2697, com uma resolução
entre amostras que variou de 1 a 10 cm. Dois dos níveis foram considerados
estéreis uma vez que apresentavam uma fraca ou quase nula quantidade
esporopolínica.
Na sondagem MD99-2331, entre 2 m e 7 m de profundidade, foi
efectuada ainda a análise semi-quantitativa de IRDs e das associações de
foraminíferos planctónicos (em 78 níveis sedimentares) assim como a
determinação do δ18O contido nas carapaças de foraminíferos planctónicos
(em 58 níveis sedimentares) e bentónicos (em 29 níveis sedimentares) com
uma resolução entre amostras que variou de 2 a 10 cm. Foram ainda
efectuadas 40 datações por
foraminíferos
planctónicos
14C
AMS (Accelerator Mass Spectrometry) em
do
tipo
Globigerina
bulloides
ou
Neogloboquadrina pachyderma (s.), ao longo dos primeiros 7 m de
sondagem.
Nos primeiros 4 m da sondagem MD03-2697 foi ainda efectuada a
análise semi-quantitativa de IRDs (em 60 níveis sedimentares) e das
associações de foraminíferos planctónicos (em 71 níveis sedimentares), a
determinação de δ18O de foraminíferos planctónicos (em 59 níveis
sedimentares) e bentónicos (em 35 níveis sedimentares) com uma resolução
entre amostras de 2 a 10 cm, e ainda 11 datações
63
14C
AMS.
F. Naughton, 2007
1. 5. 2. 2 Sondagem marinha pouco profunda VK03-58Bis
A sondagem VK03-58Bis foi colhida na plataforma continental
noroeste francesa, na Baía de Biscaia, no sector “Sud-Glénan” do complexo
silto-argiloso designado por “La Grande Vasière” (47°36’ N e 4°08’ W) a 96.8
m
de
profundidade)
usando
uma
vibrosonda,
durante
a
missão
oceanográfica Vibarmor (Fig. I.25). Esta sondagem apresenta cerca de 2.72
m de sedimento silto-argiloso e cobre aproximadamente os últimos 8 855
anos (8 855 cal yr BP). Esta elevada taxa de sedimentação permite a
obtenção de um registo climático de altíssima resolução temporal para o
Holocénico.
Nesta sondagem, foi efectuada a análise polínica em 42 níveis
sedimentares, em todos os 2 a 8 cm de sedimento ao longo dos 2.72 m de
sondagem.
Foi
ainda
efectuada
a
análise
das
associações
de
dinoflagelados e o estudo das comunidades bentónicas (nomeadamente
do gastrópode Turritella communis) em 15 amostras sedimentares e ainda 5
datações
14C
AMS nas carapaças de gastrópodes do tipo T. communis.
1. 5. 2. 3 Sondagens estuarinas
De forma a compreender o Impacto da variabilidade climática na
evolução dos sistemas costeiros do noroeste da Península Ibérica, durante a
fase final da deglaciação (Holocénico), foram utilizadas 2 sondagens, as
quais foram colhidas na zona do estuário do Douro (Fig. I.24; Fig. I.25).
As sondagens “Core 1” e “Core 1B” (situadas à distância de 50 cm
uma da outra) foram recolhidas numa zona temporariamente emersa da
Baía de S. Paio (estuário do Douro; 41°09’N; 8°38’W), com recurso a uma
sonda hidráulica (rotary perforation). A sondagem total, composta pelas
sondagens “Core 1” e “Core 1B”, é designada por sequência do Douro. A
sequência do Douro atingiu cerca de 20 m de profundidade e cobre
essencialmente o período Holocénico.
Foi efectuada a análise polínica em 9 níveis nos primeiros 7.90 m (-4.4
m OD) da sondagem“Core1”. Esta sondagem contém importantes hiatos
polínicos os quais estão associados à deposição de grandes quantidades de
material grosseiro: areia e cascalho. (OD-Ordnance Datum, representa a
profundidade corrigida em relação ao actual nível médio do mar). Na
64
F. Naughton, 2007
sondagem, “Core 1B”,
foi efectuada a análise polínica em 68 níveis
sedimentares, todos os 2 a 20 cm de sedimento entre, 10.50 a 17.40 m (-6.81
a -13.90 m OD) de profundidade. Foi ainda efectuada uma análise
sedimentológica nas sondagens “Core1 e Core 1B”, a qual compreende a
determinação de teores em carbonatos, granulometria, determinação de
teores em matéria orgânica e caracterização sedimentológica de 64
amostras, com um espaçamento entre amostras que varia de 2 a 5 cm. Foi
ainda efectuado a caracterização do nível de cascalho presente nas
sondagens estuarinas, através do estudo morfométrico de 221 balastros, e
finalmente 2 datações por
14C
AMS em material orgânico no “Core1” e 3 no
“Core 1B”.
A análise polínica de todas as sondagens referenciadas anteriormente
(MD99-2331, MD03-2697, Vk03-58Bis, “Core1” e “Core1B”) foi efectuada, no
laboratório UMR CNRS 5805 EPOC, da Universidade de Bordéus. A análise
semi-quantitativa de IRD’s e associações de foraminíferos planctónicos das
sondagens marinhas profundas, foi realizada no mesmo laboratório por
Josette Duprat. A análise de alcanonas da sondagem MD99-2331 foi
efectuada por Edouard Bard e Frauke Rostek, no laboratório CNRS UMR-6635
Cerege da Universidade de Aix-MarseilleIII. A determinação do δ18O de
foraminíferos planctónicos do tipo Globigerina bulloides, na sondagem
MD99-2331, foi executada por Bruno Malaizé, no UMR CNRS 5805 EPOC da
Universidade de Bordéus 1, enquanto que o mesmo tipo de análises, na
sondagem MD03-2697, foi efectuado por Elsa Cortijo no “Laboratoire des
Sciences du Climat et de l’Environnement (LSCE)” em Gif-sur-Yvette. Elsa
Cortijo analisou ainda, o conteúdo em δ18O dos foraminíferos bentónicos do
tipo Cibicides wuellestorfi, em ambas as sondagens marinhas profundas. A
análise radiográfica destas sondagens marinas profundas foi efectuada no
no laboratório UMR CNRS 5805 EPOC, da Universidade de Bordéus com a
ajuda de Sébastien Zaragosi e Michel Cremer.
As datações
14C
AMS destas sondagens marinhas profundas, foram
efectuadas nos laboratórios “LMC-Laboratoire de Mesure du Carbone 14”
em Saclay (França), “GifA-AMS laboratory” em Gif-sur-Yvette (França) e no
“Beta Analytic Inc.” (USA) enquanto que, as datações
65
14C
AMS da
F. Naughton, 2007
sondagem marinha pouco profunda (Vk03-58Bis) foram realizadas no
“Poznan Radiocarbon Laboratory” (Polónia). As sondagens estuarinas foram
datadas no laboratório “Beta Analytic Inc.” (USA).
A análise sedimentológica da sondagem Vk03-58Bis assim como a
contagem de T. Communis foi efectuada por Folliot no IFREMER a Brest. A
contagem de dinoflagelados foi por seu lado executada por Jean-Loius
Turon no laboratório UMR CNRS 5805 EPOC, da Universidade de Bordéus.
Uma parte da análise sedimentológica das sondagens “Core1 e Core
1B”,
nomeadamente
granulometria,
a
determinação
determinação
de
de
teores
teores
em
em
matéria
carbonatos,
orgânica
e
caracterização do nível de cascalho presente nessas sondagens estuarinas
foi efectuada pela candidata no IPIMAR-Instituto de Investigação das Pescas
e do Mar (Portugal) enquanto que, a caracterização sedimentológica da
fracção arenosa foi efectuada pela Anabela Oliveira no IPIMAR.
1. 5. 3 Cronologia e datações 14C
A principal ferramenta necessária ao estudo comparativo e à
correlação indirecta de diferentes registos paleoclimáticos do Quaternário
tardio é, sem dúvida, a cronologia absoluta.
À excepção das sequências constituídas por “varvas” (laminação
anual), a cronologia obtida em sequências sedimentares continentais e
marinhas baseia-se nas datações radiométricas.
O cálculo da datação
do
14C
14C
baseia-se no princípio de que a actividade
contido no CO2 atmosférico tem sido constante ao longo dos anos
(Libby, 1952). Contudo, vários estudos mostram que a concentração de
14C
na atmosfera variou significativamente ao longo do tempo (Linick et al.,
1986; Stuiver et al., 1991). Esta variabilidade ocorreu como resposta a
modificações da intensidade no campo geomagnético terrestre, da
actividade solar e ainda, da redistribuição de
14C
entre os diferentes tipos de
sub-sistemas climáticos, nomeadamente em função da variação da
circulação termohalina (Hughen et al., 2000). Por essas razões, torna-se
necessário converter as idades
14C
convencionais (anos BP) em idades
calibradas (anos cal BP) (BP-Before present, sendo o presente considerado
como o ano de 1950).
66
F. Naughton, 2007
Recentemente, foram elaboradas uma série de curvas de calibração
tais como: IntCal04, Marine04 e SHCal 04, as quais são recomendadas para
efectuar a calibração das idades
série
de
dados
estão
14C
convencionais (Reimer et al., 2004). A
inseridas
no
programa
CALIB
Rev
5.0
(http://calib.qub.ac.uk) (Stuiver et al., 2005).
Nas sondagens marinhas MD99-2331, MD03-2697 e Vk03-Bis, as quais se
situam no Atlântico Norte utilizou-se a série de dados marine 04.14c inserida
no programa CALIB Rev 5.0 para calibrar Idades mais jovens do que 21 786
BP (Stuiver & Reimer, 1993; Hughen et al., 2004; Stuiver et al., 2005). Nesta
calibração foi utilizado o intervalo de confiança de 2 sigma (95.4%), assim
como as áreas relativas da curva de probabilidade e, ainda, a
probabilidade mediana da distribuição de probabilidade (Telford et al.,
2004), tal como sugerido por Stuiver et al. (2005). Este programa incorpora a
correcção temporal do efeito de reservatório do oceano global, que é de
400 anos. Este valor está perfeitamente de acordo com o valor da idade
reservatório estipulada para a margem Ibérica por Bard et al. (2004).
Contudo, a utilização deste programa de calibração é limitada a idades
mais recentes do que 21 786 BP, sendo por isso necessário recorrer a outro
tipo de curvas de calibração de forma a calibrar idades
14C
convencionais
mais antigas. As datações mais antigas do que 21 786 BP foram calibradas
utilizando o polinómio simples de segunda ordem, o qual foi elaborado por
Bard et al. (2004), a partir da correlação entre curvas de variação obtidas
por diferentes indicadores paleoclimáticos marinhos e continentais, na
sondagem marinha MD95-2042 (margem Ibérica), com a curva isotópica do
oxigénio contido na sondagem de gelo GISP2:
Idade anos cal BP = -6.2724 x 10-6 x [idade14C anos BP]2 + 1.3818 x [idade14C anos BP] –1388
As datações efectuadas nas sondagens “Core1” e “Core 1B”, do
estuário do Douro, foram calibradas pela Beta Analytic Inc. (USA) utilizando o
programa INTCAL 98 (Stuiver et al., 1998).
67
F. Naughton, 2007
1. 5. 4 Indicadores paleoclimáticos
Os grãos de pólen representam a ferramenta principal utilizada no
decorrer
deste
trabalho.
Contudo,
para
além
deste
indicador
paleoclimático terrestre, as sondagens marinhas profundas foram sujeitas a
um estudo multidisciplinar, o qual inclui uma série de indicadores
paleoclimáticos marinhos e um indicador do volume de gelo acumulado nos
pólos.
1. 5. 4. 1 Variação do coberto vegetal e do clima continental
Tal como foi referido previamente, os grãos de pólen e esporos
inclusos nos sedimentos marinhos, reflectem uma imagem integral da
vegetação regional do continente adjacente (Muller, 1959; Koreneva, 1971;
Cross et al., 1966; Manten, 1966; Groot & Groot, 1966; Turon, 1984; Heusser &
Balsam, 1977; Heusser & Shackleton, 1979) e como tal são indicadores das
condições atmosféricas do continente adjacente.
De forma a identificar variações do coberto vegetal e do clima,
procedeu-se ao tratamento laboratorial, à identificação e contagem dos
esporomorfos, ao cálculo da percentagem e concentração polínica, à
elaboração de diagramas polínicos e, finalmente, à reconstrução qualitativa
e quantitativa dos parâmetros climáticos:
a) Tratamento das amostras e montagem das lâminas
Todas
as
amostras
foram
preparadas
segundo
o
protocolo
experimental descrito por de Vernal et al. (1996) ligeiramente modificado no
laboratório “Environnements et Paléoenvironnements Océaniques” (UMR
CNRS 5805 EPOC) da Universidade de Bordéus 1 (Desprat, 2005).
Foram lavados cerca de 3 a 5 cm3 de sedimento seco utilizando um
crivo de forma a recuperar a fracção inferior a 150 μm. Esta fracção final foi
decantada durante 48 horas. Eliminou-se a água em excesso com a ajuda
de uma bomba de vácuo, recuperando-se o resíduo com a ajuda de água
destilada, para um tubo de fundo cónico de 100 cm3. Este resíduo foi
centrifugado a 2500 rpm (rotações por minuto) durante 7 minutos. A água
em excesso foi retirada e colocaram-se, no tubo, duas pastilhas contendo
um número conhecido de esporos exóticos de Lycopodium (25084 grãos no
68
F. Naughton, 2007
seu total), para poder estimar posteriormente a concentração de
palinomorfos por cm3.
De forma a eliminar totalmente o conteúdo em carbonatos das
amostras, foram efectuados vários ataques químicos, a frio, utilizando ácido
clorídrico: primeiro utilizando uma solução de HCl a 10% seguido de uma
solução de HCl a 25% e posteriormente uma solução de HCl a 50%.
De seguida, as amostras foram sujeitas a dois ataques químicos com
ácido fluorídrico de forma a eliminar os silicatos contidos no sedimento. Na
primeira manipulação, foi utilizado o HF a 40% deixando-o reagir durante três
a quatro horas sobre um agitador e, na segunda, foi aplicado o HF a 70%, o
qual permaneceu em repouso, durante 48 horas.
De forma a eliminar os géis fluorissilicatados formados durante a
reacção do resíduo com o ácido fluorídrico, procedeu-se de novo a dois
ataques químicos com uma solução de HCl - 25%, a frio.
O resíduo foi lavado com água destilada a fim de eliminar todos os
excedentes deixados pelos ácidos e, finalmente, filtrado (filtro de 10 µm),
recuperando-se a fracção compreendida entre 10 µm e 150 µm (Heusser &
Stock, 1984).
O resíduo final foi montado sobre lâmina e lamela, juntamente com
algumas gotas de glicerol de forma a obter uma lâmina móvel, a qual
permite a observação dos palinomorfos em todas as suas vistas (polar e
equatorial), facilitando a sua identificação.
b) Identificação e contagem dos palinomorfos
A
contagem
dos
palinomorfos
foi
efectuada
utilizando
um
microscópio óptico Zeiss, com objectivas de X40 e X100 (óleo de imersão). A
identificação e determinação polínica foi feita com base nas características
morfológicas apresentadas pelos esporomorfos, que permitem distinguir os
diferentes taxa polínicos. Esses taxa polínicos compreendem famílias, géneros
e tipos polínicos (grupos de espécies ou géneros com características
morfológicas semelhantes). Neste último caso foi utilizado o sufixo “tipo”
(type). Esta determinação taxonómica dos pólens foi efectuada com a
ajuda de uma colecção de referência, existente no UMR CNRS 5805 EPOC
69
F. Naughton, 2007
da Universidade Bordéus 1 e de dois atlas polínicos: Moore et al. (1991) e
Reille (1992).
Nas sondagens marinhas profundas e pouco profundas e amostras
superficiais foi contado, em cada amostra, um valor mínimo de 100 grãos de
pólen (excluindo os grãos de Pinus, plantas aquáticas e os esporos), pelo
menos 20 taxa e mais de 100 grãos de Lycopodium (McAndrew & King, 1976;
Maher, 1981).
Nas sondagens e amostras superficiais estuarinas foram contados em
cada uma das amostras cerca de 300 a 350 grãos de pólen, excluindo as
plantas aquáticas e os esporos (Rull,1987), pelo menos 20 taxa e mais de 100
grãos de Lycopodium (McAndrew & King, 1976; Maher, 1981).
Os
esporomorfos
danificados
(partidos,
corroídos,
dobrados
e
escondidos) foram assinalados como indeterminados. Existem ainda outros
grãos que não foram identificados, os quais foram inseridos na categoria dos
não identificados.
c) Cálculo da percentagem e da concentração esporopolínica
A percentagem de cada taxon, também designada por frequência
polínica relativa (de acordo com Birks & Birks, 1980), foi calculada
relativamente ao somatório de base:
% taxa w = número de pólens contados do taxa w X 100/ somatório de base
Geralmente, nas sequências polínicas continentais, o somatório de
base corresponde à soma de todos os grãos de pólen de árvores, arbustos e
herbáceas, enquanto o somatório total representa a adição do somatório
de base com o somatório das plantas aquáticas, dos esporos, dos
indeterminados e dos grãos não identificados.
No entanto, os grãos de pólen do género Pinus são geralmente sobrerepresentados nos sedimentos marinhos (Heusser & Balsam, 1977; Turon, 1984)
70
F. Naughton, 2007
pelo que estes são excluídos da soma de base quando se efectua um
estudo polínico de sondagens marinhas.
Desta forma, nas sondagens marinhas profundas (MD99-2331 e MD032697) o Pinus não foi incluído no somatório de base e a sua percentagem foi
determinada em relação ao somatório total a qual inclui todos os taxa
polínicos, as plantas aquáticas, os esporos, os indeterminados e os grãos não
identificados.
A sondagem marinha pouco profunda Vk03-58Bis foi recolhida
próxima da actual linha de costa e sabendo que normalmente a sobrerepresentação de Pinus aumenta à medida que nos afastamos da linha de
costa (Muller, 1959; Groot & Groot, 1966; Bottema & Van Straaten, 1966;
Koreneva, 1966; van der Kaars & de Deckker, 2003), e tendo verificado que
este
taxa
não
aparenta
estar
sobre-representado
relativamente
às
sequências polínicas do continente adjacente, admitimos a não sobrerepresentação do mesmo na sondagem Vk03-58Bis, pelo que este taxa foi
incluído no somatório de base.
Nas sondagens estuarinas Core1 e Core1B o Pinus foi incluído também
no somatório de base.
A concentração esporopolínica contida nos sedimentos é definida
pela quantidade de esporos e pólens presentes por unidade de volume ou
de massa de sedimento. Tal concentração, é expressa em grãos/cm3 ou em
grãos/g. Com base na concentração polínica obtida em cada amostra, énos possível avaliar se as variações ocorridas, em termos de percentagens
de taxa, são reais ou se resultam apenas de efeitos estatísticos. De forma a
determinar a concentração de palinomorfos por unidade de volume
(grãos/cm3), utilizou-se a seguinte fórmula:
[taxa w] = n° de pólens contados do taxa w X [exóticos] / n° de
exóticos
d) Diagrama Polínico
A proporção de cada taxa expressa em frequências relativas para
cada nível sedimentar constitui o chamado espectro polínico.
71
F. Naughton, 2007
O diagrama polínico é a representação sucessiva e vertical de todos
os espectros polínicos, permitindo uma melhor visualização da variação
desses espectros ao longo das sequências sedimentares. Para a elaboração
dos diagramas polínicos recorreu-se ao programa PSIMPLOLL (Bennett, 1992).
Estabeleceram-se vários tipos distintos de diagramas polínicos, de
modo a compreender claramente as variações ocorridas ao longo da
sondagem, nomeadamente:
- diagramas polínicos detalhados de percentagens em função da
profundidade,
constituído
por
todos
os
taxa
polínicos,
esporos,
indeterminados e não determinados;
- diagramas sintéticos de percentagens em função da profundidade,
constituído por uma selecção dos taxa que permitem decifrar as variações
do coberto vegetal;
-diagramas sintéticos representando uma série de grupos ecológicos distintos
que permitem identificar as variações ecológicas que ocorreram ao longo
do tempo;
- diagramas de concentrações polínicas em função da profundidade.
O diagrama polínico de uma sequência sedimentar permite-nos
reconstituir associações ecológicas passadas quando comparadas com
associações actuais (princípio do uniformitarismo - “o presente é a chave do
passado”), uma vez que, na sua maioria, estas ainda persistem como
géneros actuais (Reille, 1990). A distribuição actual das diferentes formações
vegetais está directamente associada a um dado clima. Assim, conhecendo
as variações que ocorrem nas associações polínicas ao longo do tempo, énos possível deduzir através de um diagrama polínico, as variações
climáticas associadas de forma qualitativa, visto que essas associações
estão em equilíbrio com o clima.
O diagrama polínico é dividido em zonas polínicas de maneira a
facilitar a interpretação das variações do conteúdo esporopolínico em
termos de vegetação e clima. O estabelecimento de uma zona polínica
baseia-se na flutuação qualitativa de pelo menos duas curvas de taxa
72
F. Naughton, 2007
ecologicamente importantes relativamente a uma zona sub ou suprajacente
(Pons & Reille, 1986).
e) Reconstrução quantitativa de parâmetros climáticos baseada nas
associações polínicas
A reconstrução climática quantitativa de uma dada zona pode ser
efectuada em sequências polínicas utilizando a técnica dos melhores
análogos actuais (MAT-Modern Analogue Technique)(Guiot et al., 1989;
Guiot; 1990), com recurso ao programa 3Pbase (Guiot & Goeury, 1996). Esta
técnica baseia-se na selecção de 5 espectros polínicos actuais (análogos
actuais) que apresentam uma assinatura polínica semelhante àquela que
caracteriza a amostra fóssil. A base de referência polínica moderna (modern
pollen database) é constituída por cerca de 1328 espectros polínicos
provenientes da Europa, Eurásia e norte de Africa (Peyron et al., 1998; Peyron
et al., 2005), os quais incluem cerca de 103 taxa polínicos.
Cada um destes espectros polínicos actuais é representado por uma
série de parâmetros climáticos, tais como: precipitação média anual (PANNannual mean precipitation), temperatura média anual (TANN-annual mean
temperatures), temperatura média do mês mais frio (MTCO-temperature
mean of the coldest month), e a temperatura média do mês mais quente
(MTCO-temperature
mean
of
the
warmest
month), os
quais
foram
previamente interpolados utilizando uma técnica de ANN (Artificial Neural
Network) (Peyron et al., 1998).
1. 5. 4. 2 Indicadores paleoclimáticos marinhos
A associação dos vários indicadores climáticos marinhos (IRD,
associações de foraminíferos planctónicos, δ18O de de foraminíferos
planctónicos, alcanonas) efectuada nas sondagens marinhas profundas
(MD99-2331 e MD03-2697), permitiu detectar a variabilidade climática que
ocorreu no Atlântico Norte. Este estudo foi efectuado na fracção grosseira
(>150 μm) recuperada durante a lavagem das amostras utilizadas para
efectuar a análise polínica.
a) Identificação da dinâmica dos Icebergs no Atlântico Norte
73
F. Naughton, 2007
Após o colapso das grandes calotes glaciárias situadas no Hemisfério
Norte, foi libertada uma grande quantidade de icebergues os quais foram
dispersos através das correntes oceânicas. Os icebergues contêm, na sua
base, detritos de dimensão grosseira (geralmente superior a 150 μm)
designados por IRD (Ice-rafeted detritus) e, à medida que estes são
transportados para sul, fundem e libertam os seus sedimentos nos fundos
oceânicos. A contagem de detritos foi efectuada com a ajuda de uma lupa
binocular em pelo menos 10 gramas de sedimento.
b) Evolução da temperatura da massa de água superficial
A reconstrução das condições de temperatura oceânica pode ser
efectuada utilizando uma série de ferramentas distintas:
b. 1) associações de foraminíferos planctónicos
Os foraminíferos planctónicos pertencem ao Reino Protista e são
constituídos
por
uma
carapaça
mineralizada
carbonatada.
A
sua
distribuição específica segue um padrão latitudinal de variação da
temperatura da massa de água superficial, e depende das condições de
salinidade local. A presença de foraminíferos planctónicos ao longo da
coluna de água pode ser detectada até cerca de 4 000 m de
profundidade. No entanto, estes organismos são mais abundantes nos
primeiros 200 m de profundidade (próximo da zona fótica).
Foram contados e identificados pelo menos cerca de 400 indivíduos
em cada nível estudado, com base em Kennet & Srinivasan (1983). Os
foraminíferos
planctónicos
bioclimáticas
principais
Neogloboquadrina
foram
agrupados
nomeadamente:
pachyderma
em
três
polares
sinistral),
associações
(que
subpolares
inclui
(que
inclui
Neogloboquadrina pachyderma dextra, Globigerina bulloides e Turborotalia
quinqueloba)
e
tropicais
a
subtropicais
(que
inclui
ias
espécies
temperadas/frias, subtropicais, subtropicais quentes e tropicais, Globorotalia
scitula, G.inflata, G.hirsuta, G.truncatulinoides, G.crassaformis, Globigerinita
glutinata,
Globigerina
falconensis,
G.calida,
G.rubescens,
G.digitata,
Hastigerina aequilateralis, Orbulina universa, Globigerinoides ruber) (e.g. Bé,
1977; Ottens, 1991; Duprat, 1983).
74
F. Naughton, 2007
O conhecimento sobre a distribuição actual dos foraminíferos
planctónicos no oceano mundial, assim como das condições ecológicas
exigidas pelas diferentes espécies, permite-nos reconstituir as variações da
temperatura da massa de água superficial, durante o inverno (Fevereiro) e o
verão (Agosto), no passado, através da utilização de uma função de
transferência. Esta função de transferência é baseada na técnica dos
melhores análogos modernos e cuja base de dados foi estabelecida por
Pflaumann et al. (1996), e posteriormente melhorada por Elsa Cortijo
(Laboratoire des Sciences du Climat et de l’Environnement-LSCE, Gif-surYvette, França) e por Josette Duprat (UMR CNRS 5805 EPOC, Universidade de
Bordéus 1, França).
b.2) Estimativa da SST anual obtida a partir das alcanonas
O cocolitóforo Emiliania huxleyi biosintetiza alcanonas sob a forma de
uma série de componentes contendo átomos de carbono (C37-C39) os quais
são compostos por duas ou três ligações (grau de insaturação) (Volkman et
al., 1980; Volkman et al., 1995). O Índice Uk37, obtido através da equação de
Prahl (Prahl et al., 1988), permite quantificar o grau de insaturação de uma
dada série de alcanonas. Experiencias laboratoriais mostraram uma boa
correlação entre o Uk37΄ e a temperatura do crescimento da espécie E.
huxleyi permitindo a sua utilização como um indicador marinho na estimativa
de SST (Prahl et al., 1988; Rostek et al., 1993; Rosell-Melé et al.,1995). Apesar
de existirem ainda dúvidas relativamente ao impacto da sazonalidade no
sinal da SST estimado a partir deste indicador paleoclimático (Sachs et al.,
2000), iremos assumir que os resultados estimados ao longo deste trabalho
representam os valores de SST anual.
A reconstrução da SST anual foi efectuada apenas na sondagem
marinha profunda MD99-2331.
A extracção das alcanonas foi efectuada utilizando um “Dionex
Accelerated Solvent Extractor (ASE-200)” automático. Previamente aos
processos de extracção, a introdução de n-C36 na célula de ASE permite
determinar a concentração de C37. A estimativa da SST foi calculada
posteriormente utilizando a equação de Prahl et al. (1988).
75
F. Naughton, 2007
b.3) Determinação dos valores de δ18O obtidos a partir das carapaças
de foraminíferos planctónicos do tipo Globigerina bulloides
Os isótopos de oxigénio encontram-se incorporados nas moléculas de
água do mar. As moléculas de água enriquecidas em
16O
(isótopo leve) são
preferencialmente sujeitas a processos de evaporação na zona equatorial e,
posteriormente, transportadas para os pólos, deixando a água do mar
enriquecida em
18O
(isótopo pesado). Durante o seu percurso (da zona
equatorial para os pólos) ocorrem vários processos de condensação e de
precipitação durante os quais o
18O
é removido da atmosfera, chegando
aos pólos empobrecido neste isótopo pesado e enriquecido no isótopo leve.
Ao chegar aos pólos, a precipitação de neve vai enriquecer o gelo em 16O.
Sabe-se ainda que a composição isotópica do oxigénio (δ18O)
contido
nas
carapaças
de
foraminíferos
planctónicos
depende
da
temperatura, salinidade (do local) e da composição isotópica do oxigénio
da água do mar. Por esta razão, estas carapaças vão estar enriquecidas em
18O
quando as condições são favoráveis à acumulação de gelo nos pólos.
A determinação de δ18O contido nas carapaças de foraminíferos
planctónicos permite-nos determinar aproximadamente a temperatura da
massa de água superficial. Contudo, deve ter-se em consideração o facto
de que por vezes a salinidade local pode afectar ligeiramente o sinal
isotópico. Por essas razões, ao longo deste trabalho foram utilizados 2 a 3
indicadores distintos de SST tal como foi recentemente foi sugerido pelo
grupo MARGO (Multiproxy Approach for the Reconstruction of the Glacial
Ocean surface) (Kucera et al., 2005).
A análise isotópica foi efectuada em foraminíferos planctónicos do
tipo Globigerina bulloides. Os foraminíferos foram retirados da fracção
granulométrica compreendida entre 250 e 315 µm e lavados com água
destilada. A preparação de cada “aliquot” (4 a 10 indivíduos representando
uma média de 80 μg) foi efectuada com a ajuda de um ataque químico
individual no amostrador “Micromass Multiprep autosampler”. As análises
isotópicas foram efectuadas no espectrómetro “Optima Micromass mass
spectrometer” (UMR CNRS 5805 EPOC) e no “delta plus Finnigan” (LSCE). Os
resultados obtidos são expressos versus PDB.
76
F. Naughton, 2007
1. 5. 4. 3 Variações do volume de gelo acumulado nos pólos
Para além dos indicadores paleoclimáticos continentais e marinhos,
foi ainda utilizado o registo de δ18O contido nas carapaças de foraminíferos
bentónicos, o qual funciona como indicador de variações no volume de
gelo acumulado nos pólos (Shackleton, 1987).
Contrariamente ao que se passa nas massas de água de superfície,
onde as condições de temperatura variam bastante, as águas de fundo são
menos afectadas por variações de temperatura (Shackleton, 1987).
Contudo, alguns registos mostram uma grande amplitude na variabilidade
do δ18O que parece estar relacionada com variações importantes na
temperatura das massas de água profunda (Labeyrie et al., 1987; McManus
et al., 1999; Shackleton et al., 2000b). Uma forma de minimizar estes efeitos
seria estudar zonas onde as variações de temperatura de fundo aparentem
ser mínimas (Labeyrie et al., 1987).
Apesar das dificuldades inerentes às variações da temperatura de
fundo, utilizámos o δ18O contido nas carapaças de foraminíferos bentónicos
como indicador de variações do volume de gelo acumulado nos pólos.
A análise isotópica foi efectuada em foraminíferos bentónicos do tipo
Cibicides wuellerstorfi e, o processo laboratorial seguiu a metodologia
descrita para a análise isotópica dos foraminíferos planctónicos.
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Capítulo 2| Present-day and past (last 25 000 years)
marine pollen signal off western Iberia
Sinal polínico marinho presente e passado (últimos 25 000
anos) ao longo da margem oeste Ibérica
Répresentation pollinique actuelle et passée (des derniers 25
000 ans) dans les sédiments marins de la marge ibérique
occidentale
Marine Micropaleontology
in press. 2006
F. Naughton a, e , M.F. Sánchez Goñi a, b, S. Desprat a, b, J-L. Turon a, J. Duprat a, B.
Malaizé a, C. Joly a, E. Cortijo c, T. Drago d, M.C. Freitas e
a
Environnements et Paléoenvironnements Océaniques (UMR CNRS 5805 EPOC),
Université Bordeaux 1, Av. des Facultés, 33405 Talence, France
b
Ecole Pratique des Hautes Etudes, Environnements et Paléoenvironnements
Océaniques (UMR CNRS 5805 EPOC), Université Bordeaux 1, Av. des Facultés, 33405
Talence, France
c
Laboratoire des Sciences du Climat et de l’Environnement (LSCE-Vallée), Bât. 12,
avenue de la Terrasse, F-91198 Gif-sur-Yvette cedex, France
d
Centro Regional de Investigação Pesqueira do Sul , Instituto Nacional de
Investigação Agrária e Pescas (INIAP) (IPIMAR-CRIPSUL), Av. 5 de Outubro, 8700-305
Olhão, Portugal
e
Departamento e Centro de Geologia -Universidade de Lisboa, Bloco C6, 3º piso,
Campo Grande, 1749-016 Lisboa, Portugal
105
F. Naughton, 2007
Resumo
A comparação entre assinaturas polínicas actuais terrestres e
marinhas, ao longo da margem e Península Ibérica, mostra que, as
associações
polínicas
marinhas
fornecem
uma
imagem
integral
da
vegetação regional que coloniza o continente adjacente. As comunidades
florestais mediterrânicas e Atlânticas são facilmente discriminadas pelos
espectros polínicos marinhos do sul e norte, respectivamente. Os resultados
obtidos a partir das concentrações polínicas totais, juntamente com, o
conhecimento sobre a dinâmica actual das partículas sedimentares finas ao
longo da margem Ibérica, permitiu-nos estabelecer o padrão actual de
dispersão polínica, nesta região.
O
registo
paleoclimático,
relativo
aos
últimos
25 000
anos,
representado por indicadores climáticos continentais (pólen) e marinhos
(δ18O de foraminíferos planctónicos do tipo G. bulloides, Ice-rafted detritusIRD e percentages de N. pachyderma sinistrógira), ao longo de duas
sondagens (MD99-2331 e MD03-2697) colhidas na margem noroeste Ibérica,
mostra que a vegetação do noroeste da Península Ibérica respondeu
contemporaneamente à variabilidade climática detectada no Atlântico
Norte. A resposta da vegetação aos eventos de Heinrich 2 e 1 é contudo
complexa, sendo caracterizada por duas fases distintas nas médias e baixas
altitudes do noroeste Ibérico. O início de cada evento de Heinrich é
marcado, no continente, por uma forte regressão da floresta de pinheiros,
assim como pela expansão de Calluna. Este sinal é síncrono dos valores mais
pesados em δ18O e ainda da máxima expansão de N. pachyderma
sinistrógira. Isto sugere que as primeiras fases de cada evento de Heinrich
terão sido frias e húmidas, nesta região. A segunda fase de cada evento de
Heinrich é marcada pela expansão da floresta de pinheiros a qual indica a
presença de um episódio ligeiramente menos frio. Esta segunda fase é ainda
marcada
pelo
aumento
da
aridez
a
qual
é
testemunhada
pelo
desenvolvimento de plantas semi-desérticas, no continente. A comparação
do nosso registo multi-proxy com uma série de sequências polínicas
continentais e marinhas Ibéricas permitiu-nos demonstrar que o evento H1 é
o equivalente marinho do evento de Drias Antigo no continente.
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F. Naughton, 2007
A ocorrência de árvores de clima temperado durante o último
máximo glaciar (LGM- Last Glacial Maximum) e a rápida expansão de
Quercus caducifolio durante o evento de Bölling-Allerød, ao longo da nossa
sequência composta Galega, mostra que não só o sul mas também o norte
da Península Ibérica reagiram como uma zona refugio para as árvores
caducifolias durante o último período glaciário, especialmente nas médias e
baixas altitudes do Noroeste Ibérico.
Finalmente, a comparação entre as sequências terrestres e marinhas
do sul e norte da Península e margem Ibérica permitiu-nos verificar que a
vegetação respondeu ao aquecimento que caracteriza o Bölling-Allerød, ao
evento frio conhecido por Drias recente e a melhoria climática que
caracteriza o Holocénico de forma mais rápida no sul e nas médias e baixas
altitudes do noroeste da Península ibérica do que nas altas altitudes da
região nortenha como resultado da grande densidade de zonas refugio
nessas zonas durante o LGM.
Résumé
La comparaison entre le signal pollinique actuel des sédiments de la
marge Ibérique occidentale et celui de la Péninsule Ibérique montre que les
assemblages polliniques marins représentent une image intégrée de la
végétation régionale qui colonise le continent adjacent. Les communautés
forestières
présentes
sur
les
zones
biogéographiques
Ibériques
Méditerranéenne et Atlantique sont bien discriminées par les spectres
polliniques marins du Sud et du Nord de la marge Ibérique, respectivement.
L’étude de la concentration pollinique totale, ainsi que la connaissance des
modèles actuels sur la dynamique des particules fines de la marge ibérique,
nous a permis d’établir les scénarii de dispersion pollinique actuelle sur cette
région.
L’enregistrement climatique des derniers 25 000 ans, obtenu à partir
des indicateurs continentaux (pollen) et marins (δ18O des foraminifères
planctoniques
du
type
G.
bulloides,
Ice-rafted
detritus-IRD
et
les
pourcentages de N. pachyderma senestre), d’une séquence marine
composite de la marge Galicienne (MD99-2331 et MD03-2697) montre que la
végétation du Nord-ouest Ibérique à répondu de façon synchrone aux
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F. Naughton, 2007
variations climatiques détectées dans l’Atlantique Nord. La réponse de la
végétation aux événements d’Heinrich 2 et 1 (H2 et H1) est cependant
complexe et, essentiellement caractérisée par deux phases distinctes dans
les basses et moyennes altitudes du Nord-ouest Ibérique. Le début de
chaque événement d’Heinrich marqué sur la marge Ibérique par le
refroidissement des eaux de surface et par l’alourdissement du δ18O de
foraminifères planctoniques, est représenté sur le continent par une forte
régression de la forêt de pin et l’expansion des bruyères (heathers). Cela
suggère que la première phase de l’Heinrich a été froide et humide dans le
Nord-ouest de la Péninsule Ibérique. L’expansion de la forêt de pin
caractérise la deuxième phase de chacun des événements d’Heinrich,
indiquant des conditions légèrement moins froides. Pendant la deuxième
phase de l’événement H1 une augmentation de la sécheresse est indiquée
montrée
par
le
développement
des
plantes
semi-désertiques.
La
comparaison entre notre enregistrement multi-proxy de la marge Galicienne
avec des séquences polliniques Ibériques et d’autres marines de la même
marge a permis de démontrer que l’événement H1 est l’équivalent marin de
l’événement du Dryas ancien premièrement défini sur le continent.
La présence d’arbres tempérés au cours du dernier maximum glaciaire
(LGM) ainsi que l’expansion rapide du chêne caducifolié pendant
l’événement du Bolling-Allerod, déduit de notre enregistrement marin, montre
que non seulement le Sud mais aussi le Nord de la Péninsule Ibérique a été
une zone refuge d’arbres caducifoliés durant la dernière période glaciaire et,
en particulier, dans les zones de basses et moyennes altitudes.
De plus, la comparaison des enregistrements marins et continentaux
entre le Sud et le Nord permet de confirmer que la végétation a répondu au
réchauffement du Bolling-Allerod, au refroidissement correspondant au Dryas
récent et à l’amélioration climatique qui caractérise l’Holocène plus
rapidement dans le sud et dans les basses et moyennes altitudes du Nordouest Ibérique que dans les hautes altitudes de la région Nord Ibérique. Cela
indique que la densité de zones refuges pour des arbres tempérés au LGM
étais plus importante dans ces zones que dans les hautes altitudes de la
Péninsule Ibérique.
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F. Naughton, 2007
Abstract
The comparison between modern terrestrial and marine pollen signals
in and off western Iberia shows that marine pollen assemblages give an
integrated image of the regional vegetation colonising the adjacent
continent. Present-day Mediterranean and Atlantic forest communities of
Iberia are well discriminated by south and north marine pollen spectra,
respectively. Results from Total Pollen Concentration together with recognized
conceptual models of fine particle dynamics in the Iberian margin have
allowed us to establish the present-day pattern of pollen dispersion in this
region.
The 25 000 year-long record of continental (pollen) and marine (δ18O of
G. bulloides, Ice-rafted detritus-IRD and N. pachyderma s.) proxies, from the
Galician margin composite core (MD99-2331 and MD03-2697), show that
vegetation cover in north-western Iberia has responded contemporaneously
to the climate variability of the North-Atlantic. The vegetation response to the
well known North Atlantic Heinrich events 2 and 1 (H2 and H1) is however
complex and characterised by two vegetation phases at low and midatitudes of north-western Iberia. The beginning of each Heinrich event is
marked on land by an important pine forest reduction and the expansion of
heathers which are synchronous with the heaviest planktonic δ18O values and
the maxima of N. pachyderma (s.) suggesting that these first phases were
cold and wet. Pinus forest expansion characterising the second phase of
each Heinrich event indicates a less cold episode associated, during H1, with
an increase of dryness as suggested by the development of semi-desert
associations. The comparison of our Galician margin multi-proxy record with
several pollen sequences from in and off Iberia allows us to demonstrate that
H1 event is the marine equivalent of the Oldest Dryas on the continent.
The occurrence of temperate trees during the last glacial maximum
(LGM) and the rapid expansion of deciduous Quercus during the BöllingAllerød period in our Galician margin composite sequence show that not only
the southern but also north-western Iberia was a refugium zone for deciduous
trees during the last glacial period, especially at low and mid-altitude zones.
Furthermore, the comparison between southern and northern marine
and terrestrial sequences allows us to confirm that vegetation responded to
109
F. Naughton, 2007
the Bölling-Allerød warming, the Younger Dryas cold event and the Holocene
more quickly in low and mid-altitudes of north-western Iberia and in the south
than in the high altitude northern region most likely as the result of the higher
density of refugia for temperate trees in these zones during the LGM.
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F. Naughton, 2007
2. 1 Introduction
During the last decade several studies had been carried out in marine
deep-sea cores off Iberia (Hooghiemstra et al., 1992; Sánchez Goñi et al.,
1999, 2000, 2002, 2005; Boessenkool et al., 2001; Roucoux et al., 2001, 2005;
Turon et al., 2003; Tzedakis et al., 2004; Desprat, 2005; Desprat et al., 2005,
2006, in press) to understand vegetation responses to the climate variability
detected in the North Atlantic. Among these sequences, those covering the
last 25 000 years show similar vegetation changes to those recorded by the
available 25 000 year-long terrestrial records. However, no experimental
studies have been conducted in order to demonstrate that pollen grains
preserved in those marine sequences represent the regional vegetation of
the nearby continent or to understand the mechanisms involved in the
transport and dispersion of these grains from the continent to the sea. To fill
these gaps, we have compared present-day continental (including coastal
systems) pollen signatures with modern marine (including shelf and slope)
pollen assemblages. We have also determined total pollen concentration
(TPC) of those surface samples to recognize present-day patterns of pollen
dispersion in the Iberian margin.
Having assessed the reliability of the present-day pollen signal in the
upper layer sediments of MD99-2331 deep-sea core, we will compare their 25
000-year high-resolution pollen record with other marine and terrestrial pollen
sequences (Pons and Reille, 1988; Hooghiemstra et al., 1992; Peñalba, 1994;
Pérez-Obiol and Julià, 1994; Allen et al., 1996; Muñoz Sobrino et al., 1997; 2001;
2004; Peñalba et al., 1997; Von Engelbrechten, 1998; Combourieu Nebout, et
al., 1999; 2002; Sánchez Goñi and Hannon, 1999; Santos et al., 2000;
Boessenkool et al., 2001; Roucoux et al., 2001; 2005; Gil-Garcia et al., 2002;
Ruiz Zapata et al., 2002; Turon et al., 2003) to document accurately western
Iberian vegetation changes over this period. Furthermore, the direct
correlation between sea surface temperature and vegetation changes in
and off Iberia from the multiproxy study of MD99-2331 and MD03-2697 deepsea cores will allow us to link several well known terrestrial climate events with
those detected elsewhere in the North Atlantic and over Greenland.
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F. Naughton, 2007
2. 2. 1 Study area and present-day vegetation and climate
Western Iberia including Portugal and the north-western part of Spain
extends from 37°N to 43°N and comprises essentially the Minho and Sado
basins and the western part of the Douro and Tagus basins (Fig. II.1). Northwestern Spain, including the Minho basin, is influenced by the wet, relatively
cool and weakly seasonal Atlantic climate (annual precipitation mean: 9001400 mm and temperature range: -7 to 10° C) and is dominated by
deciduous Quercus forest (Q. robur, Q. pyrenaica and Q. petraea), heath
communities (Ericaceae and Calluna) and Ulex. There are also locally birch
(Betula pubescens subsp. celtiberica) and hazel (Corylus avellana) groves,
and brooms (Genista) (Alcara Ariza et al., 1987).
In the south, the Tagus and Sado basins, influenced by Mediterranean
climate (mean annual precipitation: 200-600 mm and temperature range: 4
to 14° C), are dominated by evergreen sclerophyllous forests. Q. rotundifolia
and Q. suber forests with Phillyrea angustifolia and Pistacia terebinthus
colonise the western basins while Q. rotundifolia and Q. coccifera woodlands
associated with Juniperus communis and Pinus halepensis occupies the
eastern part. In the warmest zones, thermophilous elements such as Pistacia
lentiscus and Olea sylvestris form the forests. Middle altitudes (700-1000 m
a.s.l.) are dominated by deciduous Quercus forest (Q. pyrenaica and Q.
faginea) associated with northern European species such as Taxus baccata.
The degradation of this forest produces two types of brush communities:
rockrose shrublands (Cistaceae) in zones with precipitation between 600 and
1000 mm and heath communities (Ericaceae) in wetter zones.
Between both regions, there is a transitional zone which includes the
hydrographic basin of the Douro. This zone is characterised by high
precipitation values (700 to 1000 mm/year) and winter temperatures between
4 and -4°C. At high altitudes, the wettest and coldest zones reach 1600
mm/year and -8°C, respectively (Polunin and Walters, 1985). The oceanic
influence is particularly important in the northwest of the basin, where the Q.
robur and Q. suber association predominates (Braun-Blanquet et al., 1956).
The spread of both Pinus pinaster and Eucalyptus globulus has been favoured
112
F. Naughton, 2007
by anthropic impact. The understory vegetation is largely dominated by Ulex,
in association with heaths. The river margins are colonized by Alnus glutinosa,
Fraxinus angustifolia, Ulmus spp., Salix spp. and Populus spp..
Fig. II.1 | Fig. 1- Study area. Dashed line divides the Atlantic and Mediterranean biogeographical zones
(Blanco Castro et al., 1997). White circles with a dark point represent the top samples analysed in this study;
white circles represent the modern samples from the European Pollen Database; white circles with a cross
represent the studied cores sites (MD03-2697 and MD99-2331); dark circles represent marine and terrestrial
core sites used for comparison with our study. Continental sequences: a) Square A locates sequences 1 to
5: 1- Laguna de la Roya (Allen et al., 1996), 2- Sanabria March (Allen et al., 1996), 3-Laguna de las
Sanguijuelas (Muñoz Sobrino et al., 2004), 4-Lleguna (Muñoz Sobrino et al., 2004), 5- Pozo do Carballal
(Muñoz Sobrino et al., 1997); b) Sites 6 to 13 correspond to: 6- Laguna Lucenza (Santos et al., 2000); 7Lagoa Lucenza (Muñoz Sobrino et al., 2001); 8-Lago de Ajo (Allen et al., 1996); 9- Los Tornos (Peñalba,
1994); 10-Saldropo (Peñalba, 1994); 11-Belate (Peñalba, 1994); 12- Atxuri (Peñalba, 1994); 13- Banyoles
(Pérez-Obiol and Julià, 1994); c) Square B includes sequences 14 to 19: 14- Quintanar de la Sierra (Peñalba
et al., 1997); 15- Sierra de Neila - Quintanar de la Sierra (Ruiz Zapata et al., 2002); 16- Hoyos de Iregua (GilGarcia et al., 2002); 17- Laguna Masegosa (Von Engelbrechten, 1998); 18- Laguna Negra (Von
Engelbrechten, 1998); 19- Las Pardillas lake (Sánchez Goñi and Hannon, 1999); 20- Padul (Pons and Reille,
1988); 21- Mougás (Gómez-Orellana et al., 1998); 22- Charco da Candieira (Van der Knaap and Van
Leeuwen, 1995). The marine cores represented on the map are: 8057 B (Hooghiemstra et al., 1992), SO756KL (Boessenkool et al., 2001), SU81-18 (Turon et al., 2003) and ODP 976 (Combourieu Nebout, et al., 1999;
2002) and MD95-2039 (Roucoux et al., 2001; 2005).
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F. Naughton, 2007
2. 2. 2 Oceanography
The western Iberian margin is dominated by the surface Portugal
Current system (PCS) which is composed of the slow equatorward current in
the open ocean (Arhan et al., 1994) and the fast, seasonally reversing coastal
current (Ambar and Fiúza, 1994; Barton, 1998) (Fig. II.2). During the summer,
the Azores high pressure cell is located in the central North Atlantic and the
Greenland low is weak. This situation generates northerly and northwesterly
prevailing winds (Fig. II.1) which favour the occurrence of upwelling events
and a southward surface circulation (Fiúza et al., 1982; Haynes and Barton,
1990) near the shelf break in the upper 50-100 m (Álvarez-Salgado et al.,
2003). The resultant upwelled cold and nutrient-rich Eastern North Atlantic
Central Water of subpolar sources (ENACWsp) is transported northward of 45°
N. Warm, salty and nutrient-poor Eastern North Atlantic Central Water of
subtropical origin (ENACWst) is transported to the south of 40° N (Fiúza, 1984;
Rios et al., 1992) (Fig. II.2).
Fig. II.2 | West to east scheme of the different water masses from the western Iberian margin (adapted from
Sprangers et al., 2004). White circles with a dark point represent southward water flow and white circle with
a cross represent northward water flow. PCS-Portugal Current System; ENACW st-Eastern North Atlantic
Central Water of subtropical origin; ENACW sp-Eastern North Atlantic Central Water of subpolar origin; MSWMediterranean Sea Water; LSW-Labrador Sea Water; NADW-North Atlantic Deep Water.
During the winter the Azores high pressure cell is located off the
northwest African coast and the Greenland low is deep and situated off
south-eastern Greenland. The pressure gradient between the two systems
results in an onshore and slightly northward wind off Iberia (Fig. II.1) triggering
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F. Naughton, 2007
downwelling processes and a northward surface circulation (Frouin et al.,
1990; Haynes and Barton, 1990). This reversion of hydrological paths starts in
the end of summer in September-October and it persists until March-April
representing the well known Portugal Coastal Counter Current (PCCC)
(Ambar et al., 1986). This poleward flow is narrow (30 km wide) and it
transports warm and salty waters (ENACWst) in the upper 200-300 m to the
North (Pingree and Le Cann, 1990).
Below the Central Waters system, between 550 m and 1500 m depth, the
Mediterranean Sea Water (MSW) consisting of high salinity and relatively
warm water mass is transported northward (Mazé et al., 1997) (Fig. II.2).
However, the salinity of the MSW decreases highly at latitudes higher than 41°
N by mixing with the underlying low-salininity Labrador Sea water (LSW)
(McCave and Hall, 2002). This LSW is one of the three water masses included
in the North Atlantic Deep Water (NADW) over the western Iberian margin
(Huthnance et al., 2002).
2. 2. 3 Morphology and recent sedimentation
The Iberian margin is characterised by a relatively narrow shelf (30-50
km wide) with a steep irregular slope plunging to the oceanic abyssal plain
(Fig. II.3a). This margin is cut off by deep canyons like Mugia, Porto, Aveiro,
Nazaré, Cascais, Lisbon, Setúbal and S.Vicente. The largest canyons (Nazaré,
Setúbal) dissect the entire continental shelf, capturing sediments carried over
the shelf and upper slope by alongshore currents, providing a direct conduit
of particles from the upper shelf to the deep-sea (Vanney and Mougenot,
1981). Some canyons, e.g. Setúbal, start close by the present-day coastline
and have a direct connection to the river mouth, while others, such as the
Porto Canyon, begin only at the shelf edge and play a minor role in the
interception of shelf material at the present-day sea level. All Iberian canyons
were probably more active during the period of low sea-level (Van Weering
and McCave, 2002). The lower and upper slopes are also intersected by
several seamounts as Vigo (VS), Vasco da Gama (VDGS), Porto (PS), Tore (TS),
by the Galicia Bank and several tectonic depressions (Vanney and
Mougenot, 1981).
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F. Naughton, 2007
Fig. II.3 | a) Morphology of the Iberian margin. Location of the surface samples from b) north-western
Iberian margin and c) south-western Iberian margin. White arrows indicate the present-day pattern of
pollen dispersion in the western Iberian margin.
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F. Naughton, 2007
2. 2. 3. 1 North-western Iberian margin
In north-western Iberia, five rivers (Douro, Ave, Cávado, Lima and
Minho) release large amounts of sediments to the adjacent continental
margin. The Douro is the main sediment supplier to the adjacent shelf (~8.2 x
109m3 annual mean discharge) followed by the Minho river (Dias et al., 2002;
Jouanneau et al., 2002; Oliveira et al., 2002) (Fig. II.3b). They are 927 km and
300 km long, draining a catchment area of 97 700 km2 and 17 100 km2,
respectively (Loureiro et al., 1986). Above 42° N, rivers are replaced by rias
(Vigo, Pontevedra, Arousa and Muros), which act essentially as sediment
traps, preventing particle input to the adjacent margin (Dias et al., 2002;
Jouanneau et al., 2002).
The northern Portuguese continental shelf is composed of a) an inner
shelf zone (<30 m depth) with fine and well sorted sands, b) a mid-shelf zone
of coarse sands and gravels, and c) a carbonate-rich outer shelf zone with
medium sand (Van Weering et al., 2002). Within the shelf, there are two mud
patches (Douro and Galicia) located offshore from the river inlets separated
by a mud free zone (Lopez-Jamar et al., 1992) (Fig. II.3b). The mud patch
growth depends on the sediment supply, morphological barriers and
hydrological conditions (Dias et al., 2002; Jouanneau et al., 2002).
Sedimentation on the north-western Iberian margin is complex and
essentially sustained by episodic flood events (Dias et al., 2002) and/or during
maximal episodes of river outflow (Araújo et al., 1994; Drago et al., 1998). Fine
sediments, after being released by rivers, are transported in nepheloid layers
(Bottom- BNL, intermediate-INL and surface-SNL) to the outer shelf. Oliveira et
al. (1999) have shown a seaward decrease of sediment concentrations in all
nepheloid layers and that currents and waves induce resuspension of bottom
sediments from Douro and Minho muddy deposits, especially during extreme
storm events. During these extreme events, such as southwesterly storms
(downwelling conditions), the spread of BNLs might be blocked by rocky
outcrops (Drago et al., 1999) that operate as a barrier to cross-shelf transfers
(Jouanneau et al., 2002; Van Weering et al., 2002) stimulating a poleward
sediment transport (Drago et al., 1998; Dias et al., 2002; Jouanneau et al.,
2002, Van Weering et al., 2002). Sporadically, waves associated with those
storms are able to induce resuspension of fine deposits spreading offshore the
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F. Naughton, 2007
BNL (Vitorino et al., 2002) nourishing the INL (Oliveira et al. 2002). These
extreme events contribute to an important export of fine sediments (Vitorino
et al., 2002) and occasionally of coarse fraction (Dias, 1987) to the upper
slope. Current reversals, probably caused by the presence of local slope
eddies, can also allow some down-slope transport of particles (Pingree and
LeCann, 1992). During upwelling conditions, fine sediment export is restricted
to the shelf edge (McCave and Hall 2002; Van Weering et al., 2002).
However, lateral sediment exchange can be favoured by offshore filaments
stretching westward (Huthnance et al., 2002).
MD99-2331 and MD03-2697 twin deep-sea cores, located northwestern of the Mesozoic and Cenozoic outcrops, mostly receive sediments
coming from the Douro and Minho rivers, especially during downwelling
conditions.
2. 2. 3. 2 South-western Iberian margin
In the south-western Iberian margin, the Tagus river is the primary
sediment supplier followed by Sado river to the shore (Dias, 1987; Jouanneau
et al., 1998) (Fig. II.3a and Fig. II.3c). The Tagus river is 1110 km long draining a
catchment area of 80 600 km2 with 400 m3. s-1 of annual mean flow (Vale,
1990). The Sado river is 175 km long, drains a catchment area of 7 640 km2
and yields less than 10 m3. s-1 of annual mean discharge (Loureiro et al., 1986).
Differences between both river discharge and littoral currents influence the
sediment distribution along the shelf (Jouanneau et al., 1998). The mud patch
is located offshore of Tagus river basin and covers the entire continental shelf
(Araújo et al., 2002). During summer, suspended particulate matter (SPM)
concentration in the mouth of the Tagus estuary is four times higher than that
of the Sado, and the nepheloid layer can extend 30 km westward
(Jouanneau et al., 1998). Fine sediments are essentially exported to the slope
and adjacent abyssal plains through the canyons of Cascais, Lisbon and
Setúbal (Jouanneau et al., 1998) and by offshore filaments (Huthnance et al.,
2002).
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F. Naughton, 2007
2. 3. 1 Deep-sea cores: MD99-2331 and MD03-2697
MD99-2331 (42° 09’ 00 N, 09° 40’ 90 W; 2110 m depth) and MD03-2697
(42° 09’ 59 N, 59° 42’ 10 W; 2164 m depth) deep-sea cores were retrieved in
the Galician margin (north-west of Iberia) using a CALYPSO corer during the
GINNA (IMAGES V) and PICABIA oceanographic cruises on board the R/V
Marion Dufresne (Fig. II.1). MD99-2331 and MD03-2697 are 37.2 m and 41.23 m
long, respectively, covering Marine Isotopic Stages (MIS) 1 to 11. X-ray analysis
using SCOPIX image-processing (Migeon et al., 1999) has shown a well
preserved sedimentary sequence in core MD03-2697 while core MD99-2331
sees a sediment mixing zone between 1.10 m and 1.90 m of core depth. In
order to obtain a detailed palaeoclimatic sequence for the last 25 000 years
in and off NW Iberia, we have built a composite record assembling the MIS 1
interval of core MD03-2697 with the MIS 2 interval of core MD99-2331.
2. 3. 1. 1 Radiometric dating
Seven levels of MD03-2697 and twenty levels from MD99-2331 were
dated by AMS
14C
on Globigerina bulloides and Neogloboquadrina
pachyderma (s.) at Beta Analytic Inc (Beta), at Gif-sur-Yvette (Gif) and at
Laboratoire de Mesure du Carbone 14-Saclay (LMC), indicating that this
sequence covers the last 25 000 years (Tab. II.1). All radiocarbon dates were
corrected for marine age reservoir difference (400 years) (Bard et al., 2004).
The samples presenting conventional AMS
14C
younger than 21 786 BP
were calibrated by using CALIB Rev 5.0 program and "global" marine
calibration dataset (marine 04.14c) (Stuiver and Reimer, 1993; Hughen et al.,
2004; Stuiver et al., 2005). We use 95.4% (2 sigma) confidence intervals and
their relative areas under the probability curve as well as the median
probability of the probability distribution (Telford et al., 2004).
14C
radiometric
ages older than 21 786 yr BP were calibrated by matching the obtained
conventional AMS 14C with the calendar ages estimated for MD95-2042 deepsea core by Bard et al. (2004).
119
F. Naughton, 2007
In this paper, we will use
14C
ages (yr BP) corrected for the marine
reservoir effect (of 400 years) instead of calibrated ages (cal yr BP) because
in most of the Iberian terrestrial sequences, calendar ages are not available.
Lab
code
Core
depth
(cm)
Beta
2131134
Beta
2131135
Beta
003257
Beta
2131136
Beta
2131137
Beta
003258
Beta
003259
LMC14
001231
LMC14
001232
GIF
102377
LMC14
001233
LMC14
001235
LMC14
001236
LMC14
001237
LMC14
002445
GIF
101109
GIF
102373
LMC14
002446
LMC14
001845
LMC14
001846
LMC14
001847
LMC14
001849
LMC14
001850
GIF
102378
LMC14
001851
LMC14
001852
LMC14
001853
MD03 2697
20
MD03 2697
40
MD03 2697
70
MD03 2697
80
MD03 2697
110
MD03 2697
150
MD03 2697
200
MD99 2331
200
MD99 2331
205
MD99 2331
220
MD99 2331
222
MD99 2331
228
MD99 2331
235
MD99 2331
242
MD99 2331
260
MD99 2331
290
MD99 2331
570
MD99 2331
590
MD99 2331
595
MD99 2331
600
MD99 2331
607
MD99 2331
620
MD99 2331
623
MD99 2331
630
MD99 2331
637
MD99 2331
650
MD99 2331
655
Material
Conv.
AMS 14C
age BP
Conv.
AMS 14C
age BP
G. bulloides
2880
G. bulloides
error
95.4 % (2σ)
Cal BP
age ranges
Cal BP age
median
probability
2480
40
2501:2739
2656
4760
4360
40
4866:5198
5008
G. bulloides
7435
7035
50
7783:7998
7895
G. bulloides
7470
7070
40
7835:8014
7930
G. bulloides
9940
9540
40
10705:11084
10896
G. bulloides
11920
11520
60
13233:13486
13353
G. bulloides
12520
12120
60
13816:14111
13965
13640
13240
80
15303:16099
15679
13810
13410
80
15524:16359
15922
14130
13730
120
15898:16828
16342
13920
13520
90
13930
13530
80
15130
14730
15060
G. bulloides
N.
pachyderma
N.
pachyderma
N.
pachyderma
N.
pachyderma
N.
pachyderma
N.
pachyderma
N.
pachyderma
(-400 yr)
15658:16520
a
15686:16521
16067a
a
16081a
90
17250:18182
17848
14660
90
17170:18038
17722
15540
15140
90
18405:18723
18520
G. bulloides
16170
15770
130
18787:19265
18983
G. bulloides
19770
19370
170
22534:23622
23038
G. bulloides
22290
21890
170
~ 25950b,c
~ 25950b,c
G. bulloides
20860
20460
250
23931:25369
24542
G. bulloides
20550
20150
240
23450:24803
24119
G. bulloides
20460
20060
140
23652:24405
24016
G. bulloides
21620
21220
160
25301:26000
25626
G. bulloides
21740
21340
160
25439:26000
25730
N.
pachyderma
22690
22290
180
~ 26350b,c
~ 26350b,c
G. bulloides
22150
21750
170
~ 25800c
~ 25800c
G. bulloides
22440
22040
170
~ 26000c
~ 26000c
G. bulloides
22430
22030
180
~ 26000c
~ 26000c
Tab. II.1| Radiocarbon ages of MD99-2331 and MD03-2697 deep-sea cores. a Not acceptable dating
(bioturbated layers); b Radiocarbon dates too old (not used); c dates calibrated by matching conventional
AMS 14C with calendar ages estimated for MD95-2042 deep-sea core by Bard et al. (2004).
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F. Naughton, 2007
2. 3. 1. 2 Marine proxy analyses
The planktonic isotopic record of MD99-2331 covering MIS 6 to MIS 1
has been published in Gouzy et al. (2004). However, additional stable isotope
measurements of planktonic foraminifera have been done to refine the
planktonic isotopic record of the MIS 2 interval.
In total, 56 measurements have been made at 2 to 10 cm sample
resolution. For MIS 1, 29 levels with a sample spacing of 5 to 10 cm have been
analysed in the MD03-2697 sequence. These measurements have been
carried out on the 250–315 µm fraction of Globigerina bulloides previously
cleaned with distilled water. Each aliquot, including 8–10 specimens and
representing a mean weight of 80 μg, was prepared in the Micromass
Multiprep autosampler, using an individual acid attack for each sample.
The CO2 gas extracted has been analyzed against NBS 19 standard,
taken as an international reference standard.
Isotopic analysis of MD99-2331 has been carried out using an Optima
Micromass mass spectrometer in the UMR CNRS 5805 EPOC (Environnements
et Paléoenvironnements Océaniques) at Bordeaux 1 University and, those of
MD03-2697 were performed using a delta plus Finnigan at the LSCE
(Laboratoire des Sciences du Climat et de l’Environnement). The mean
external reproducibility of powdered carbonate standards is ±0.05‰ for
oxygen. Results are presented versus PDB.
Polar foraminifera, N. pachyderma (s.), counting include 79 levels (2 to
10 cm of sample spacing) and 40 levels (5 to 10 cm of sample spacing) from
MD99-2331 and MD03-2697, respectively.
IRD semiquantitative analysis as been carried out in 78 levels (2 to 10
cm sample spacing) and 30 levels (5 to 10 cm sample spacing) from MD992331 and MD03-2697, respectively. In this study, only the total concentrations
of the lithic grains were considered.
Both analyses were performed on the >150 μm sand-size fraction which
was obtained according to classic sedimentological procedure.
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F. Naughton, 2007
2. 3. 1. 3 Pollen analysis
110 and 22 samples with a sample spacing of 2 to 10 cm and 5 to 10
cm were analysed from MD99-2331 and MD03-2697, respectively. In each 1
cm-thickness sample, 3 to 5 cm3 of sediment were treated for pollen analysis.
The treatment of the samples from both deep-sea cores (MD99-2331
and MD03-2697) followed the procedure described by de Vernal et al. (1996),
slightly modified at the UMR CNRS 5805 EPOC (Desprat, 2005).
Palynological treatment consists of pollen concentration by chemical
digestion using cold HCl (at 10%, 25% and 50%) and cold HF (at 40% and 70%)
to eliminate carbonates and silicates, respectively. A Lycopodium spike of
known concentration has been added to each sample to calculate total
pollen (including spores) concentrations. The residue was sieved through a 10
µm nylon mesh screen (Heusser and Stock, 1984) and mounted in bidistillate
glycerine.
A Zeiss microscope with x550 and x1250
(oil immersion)
magnifications was used for pollen observation and counting.
Pollen identifications were achieved via comparison with specialised
atlases (Moore et al., 1991; Reille, 1992) together with the pollen reference
collection of the UMR CNRS 5805 EPOC. At least 100 pollen grains (excluding
Pinus, aquatic plants and spores) and 20 pollen types were counted in each
of the 142 samples (deep-sea cores and modern samples, cf. section 2. 3. 2)
analysed to obtain statistically reliable pollen spectra (McAndrew and King,
1976).
Pollen percentages were calculated based on the main pollen sum
which excludes aquatic plants, spores, indeterminate and unknown pollen
grains. Because Pinus grains are usually over-represented in marine sediments
(Heusser and Balsam, 1977), they are also excluded from the main sum and
their percentages are determined by using the total sum (pollen + spores +
indeterminable + unknowns).
2. 3. 2 Modern pollen samples
We have analysed the pollen grains of 10 top samples from several
estuarine, shelf and marine sedimentary sequences retrieved in and off
western Iberia (Fig. II.1, Tab. II.2).
122
F. Naughton, 2007
The high percentages of Pinus detected in these top samples confirm
that they represent the last 0-350 yr, since it is well known that Pinus
reforestation in western Iberia started in the seventeenth century (Valdès and
Gil Sanchez, 2001). Because major vegetation changes are not detected in
percentage pollen diagrams for the last centuries in this region (Desprat et al.,
2003), we assume that our modern samples represent present-day pollen
signatures.
The resulting marine and coastal modern pollen assemblages have
been compared with 12 terrestrial pollen samples including moss samples,
surface sediments and top of peat bog and lake sequences, of both the
Mediterranean and Atlantic parts of western Iberia stored in the European
Pollen Database, http:/www.imep-cnrs.com/pages/EPD.htm, (Peyron et al.,
1998; Barboni et al., 2004) (Fig. II.1).
Sample name
MD95-2042
1FP8-1
MD99-2332
Barreiro
MD04-2814 CQ
Laquasup
Depth
(cm)
Top
(0-1)
Top
(0-1)
Top
(0-1)
Top
(0-1)
Top
(0-1)
Latitude
Longitude
Water depth
(m)
Year of
sampling
37°48’N
10°10’W
3148
1995
38°01’N
09°20’W
980
2003
38°33’N
09°22’W
97
1999
38°40’N
09°07’W
0
1999
40°37’N
09°52’W
2449
2004
41°09’N
08°38’W
0
2001
41°09’N
09°01’W
81
2002
41°48’N
09°04’W
107
1992
42°09’N
09°42’W
2120
1999
42°14’N
08°47’W
45
1990
Top
(0-5)
Top
Po 287-13-2G
CG11
MD99-2331
(0-1)
Top
(0-1)
3-4
Top
Vir-18
(0-1)
Tab. II.2| Location, water depth and year of sample sampling from coastal, shelf and slope sequences of
the Iberian margin.
123
F. Naughton, 2007
2. 4. 1 Present day pollen signature
2. 4. 1. 1 Western Iberian terrestrial sites
Figure II.4 shows pollen spectra from several modern samples collected
in western Iberian Península. Pollen assemblages from surface samples
located above 42°N (COV1, MUN1, MUN3, ES09, E258 and ES62), record the
Atlantic deciduous forest (Figs. II.1 and II.4). However, the dominant tree
species differs from place to place reflecting the heterogeneity of the
vegetation cover of this region (Fig. II.4). For example, deciduous Quercus is
the most important tree pollen in samples ES09, ES258, MUN3 and ES62 while
Corylus dominates COV1 pollen spectrum and Betula that of MUN1. The
pollen signal from the southern samples (FRA1, FRA4, GAT1, EXT1 and EXT2)
represents Mediterranean plant communities, essentially composed of
evergreen Quercus (Quercus ilex-type) and Olea (Figs. II.1 and II.4). However,
FRA4 sample also includes relatively high percentages of pollen of deciduous
trees, similar to those found in pollen spectra from north-western Iberia. This
sample, though located in the Mediterranean region, comes from a high
altitude deciduous oak forest zone. Within this southern region, sample ESO6
collected in the coastal area reflects open vegetation resulting of saline
conditions and sandy soils, preventing the development of deciduous and
perennial forests. These southern samples also reflect the mosaic of the
vegetation colonising present day southern Iberia.
We notice that, the southern pollen samples show higher percentages
of Mediterranean plants than the north-western Iberian samples (Fig. II.4),
clearly discriminating between Mediterranean and Atlantic plant community
sources, respectively.
124
F. Naughton, 2007
Fig. II.4 | Pollen spectra from western Iberian modern samples. Total temperate and humid (Tot.
Temp./Hum.) trees includes: Alnus, Betula, Corylus, deciduous Quercus and other temperate and humid
species (Acer, Fagus, Fraxinus, Salix, Tilia, Ulmus, Hedera helix, Myrica and Vitis). Total mediterranean (Tot.
Mediter.) plants includes: evergreen Quercus, Olea and Cistus. Taraxacum-type, Asteraceae, Poaceae,
Ericaceae and Calluna represent the ubiquist group. Semi-desert plants include Ephedra,
Chenopodiaceae and Artemisia. Climate parameters: Alt: Altitude; PP: Precipitation; MTCO: Mean
temperature of the coldest month; MTWA: Mean temperature of the warmest month; TANN: Annual
temperature.
125
F. Naughton, 2007
2. 4. 1. 2 Western Iberian estuarine and margin sites
Estuarine pollen samples VIR-18 (Ría de Vigo) and Laquasup (Douro
estuary), are marked by relatively high percentages of deciduous forest
reflecting the present-day vegetation of north-western Iberia (Figs. II.5 and
II.1). Shelf and slope samples (MD99-2331, CG11, Po 287-13-2G and MD042814 CQ), located in the adjacent margin, reproduce the same pollen signal
as that of these northern estuarine samples. Southern samples from estuarine
(Barreiro) and margin (MD99-2332, FP8-1 and MD95-2042) sites present in turn
higher percentages of Mediterranean plants than northern sites (Fig. II.5). As
in terrestrial samples, Mediterranean and Atlantic plant communities are well
discriminated in the pollen signal from estuarine and margin sites (Figs. II.5
and II.1). It is important to note that the estuarine pollen assemblages are
more similar to the marine ones than to the terrestrial pollen spectra. Indeed,
estuarine sediments contain pollen from the regional vegetation which
colonises the hydrographic basin while terrestrial samples mainly reflect local
vegetation (Figs. II.4 and II.5). This indicates, as previous studies have already
shown for the south-western French margin (Turon, 1984), that pollen spectra
off north-western Iberia reflect an integrated image of the regional
vegetation of the adjacent continent. Pinus percentages from western
Iberian samples are relatively low when compared with estuarine, shelf and
slope samples off this region (Fig. II.5). This is in agreement with the observed
overrepresentation of Pinus pollen in marine sediments in general and, in
particular, off south-western Europe (Turon, 1984). Other works have further
shown that Pinus pollen percentages increase seawards (Heusser and Balsam,
1977; Heusser and Shackleton, 1979). Our study confirms this pattern in the
north and south-western margins. However, MD95-2042 site, representing the
farthest sample from the coast line, presents weaker percentages of Pinus
pollen than estuarine and the other marine samples.
Despite the general Pinus overrepresentation in marine sediments, the
good correlation between both terrestrial and marine present-day pollen
signatures in and off western Iberia confirm the reliability of past vegetation
and climate change reconstructions of this region proposed by previous
works on western Iberian margin cores (Hooghiemstra et al., 1992; Sánchez
Goñi et al., 1999, 2000, 2002, 2005; Boessenkool et al., 2001; Roucoux et al.,
126
F. Naughton, 2007
2001, 2005; Turon et al., 2003; Tzedakis et al., 2004; Desprat et al., 2005, 2006, in
press).
Fig. II.5 | Pollen assemblages of top samples from coastal and marine western Iberian sites (see also
caption of Fig. II.4). TPC: total pollen concentration.
127
F. Naughton, 2007
2. 4. 2 Present-day pollen transport patterns
Previous works on coastal zones with complex fluvial systems have
shown that pollen is mainly transported to the sea by rivers and streams
(Muller 1959; Bottema and Van Straaten, 1966; Peck, 1973; Heusser and
Balsam, 1977). The western Iberian margin, close to several important
hydrographic basins such as Tagus and Sado in the south and Douro and
Minho in the north, mainly receives pollen through fluvial transport (Fig. II.3).
Furthermore, north-western prevailing winds in both north and southern
regions probably impede substantial direct airborne transport of pollen
seaward. This pattern of fluvial transport contrasts with others, e.g. northwestern Africa, associated with an arid environment, where pollen grains are
mainly seaward transported by the wind (Dupont et al., 2000; Hooghiemstra
et al., 2006). Indeed, the distribution of the TPC shows (Fig. II.5) that the
highest TPC values are found in samples from coastal areas such as the Douro
estuary (Laquasup: 44,399x103 grains/cm3) and the Ría de Vigo (VIR-18:
65,443x103 grains/cm3). Barreiro sample is an exceptional case with low TPC
(18 x103 grains/cm3) probably because it was collected far away from the
Tagus main channel and likely receiving pollen only from the local
vegetation. Shelf surface samples present intermediate concentration values
(CG11: 30,468x103 grains/cm3, Po 287-13-2G: 42, 600x103 grains/cm3 and
MD99-32b: 56,396x103 grains/cm3) and finally slope samples attain the lowest
TPC values (MD99-2331: 1,924x103grains/cm3, MD04-2814 CQ: 1,886x103
grains/cm3, MD95-2042: 2,153x103 grains/cm3 and IFP8: 3,489x103 grains/cm3).
Our
work
shows
that
a
seaward
decrease
of
total
pollen
concentrations occurs on the Iberian margin following the estuary-shelf-slope
transect (Fig. II.3). This pattern coincides with that observed in other margin
zones around the world showing a seaward decrease in total pollen
concentration with maximum values close to the mouth of the river systems
(Muller, 1959; Bottema and Van Straaten, 1966; Cross et al., 1966; Groot and
Groot, 1966, 1971; Koreneva, 1966; Stanley, 1966; Mudie, 1982; Turon, 1984;
Van der Kaars and de Deckker, 2003).
Based on several studies of sedimentary dynamics on the north-western
Iberian margin (Araújo et al., 1994; Drago et al., 1998; Dias et al., 2002;
Huthnance et al., 2002; Jouanneau et al., 2002; Oliveira et al., 2002; Van
128
F. Naughton, 2007
Weering et al., 2002; Vitorino et al., 2002), we propose a pattern of pollen
dispersion for this region (Fig. II.3b). This pattern is similar to the distribution
model of fine terrigenous particles proposed by Dias et al. (2002). Pollen and
spores, once immersed behave in a similar manner to fine sedimentary
particles (Chmura and Eisma, 1995). After being released by rivers (mainly
Douro followed by Minho), pollen grains, are enclosed in nepheloid layers
and transported to the shelf until getting blocked by the rocky outcrops. In
winter, during downwelling conditions pollen grains are then transported
polewards, firstly deposited in the Douro mud patch (S-N direction) then in the
Galicia mud patch, and finally they flow westward to the deep-sea. Only
small quantities of pollen grains can be transported directly to the outer shelf
and upper slope under extreme stormy events. In summer, under upwelling
conditions, pollen transfer to the slope must be restricted to offshore filaments
as suggested by Huthnance et al. (2002) for the fine sediments.
In the southern Iberian margin, TPC values also decrease seawards as
in the northern region (Fig. II.5). Our study suggests that pollen grains released
by the Tagus and to a lesser extent by the Sado river, are partially deposited
in the shelf and transported to the south and seaward by littoral and oceanic
currents probably during upwelling conditions (Fig. II.3c). Pollen grains are
probably transported by the southern canyons from the shelf to the slope and
abyssal plain following the fine particle pathway suggested by several works
on sedimentary dynamics in this region (Dias, 1987; Jouanneau et al., 1998;
Araújo et al., 2002).
2. 4. 3 Climatic and vegetational response in western Iberia to North
Atlantic climatic events over the last 25 000 years
The comparison of the high resolution pollen composite record from
the Galician margin (Fig. II.6; Tab. II.3), with other marine and terrestrial pollen
sequences (Figs. II.1 and II.7; Tabs. II.4 and II.5) document the vegetation
changes that occurred in the Iberian Península over the last 25 000 years.
Moreover, the direct correlation between marine proxies and vegetation
changes from this record will allow us to accurately evaluate the vegetation
response to the climatic events detected elsewhere in the North Atlantic
Ocean and over Greenland.
129
F. Naughton, 2007
Fig. II.6 | Galician margin composite record (MD99-2331 and MD03-2697 deep-sea cores). From the left to
the right: corrected radiocarbon ages; marine proxies: δ18O of G. bulloides, % N. pachyderma (s.), icerafted detritus (IRD), Marine and Greenland climatic events; % pollen taxa; pollen zones and
chronostratigraphy. Pollen zones were established using qualitative and quantitative fluctuations of a
minimum of 2 curves of ecologically important taxa (Pons and Reille, 1986). They are defined by the
abbreviated name of the core (MD31 or MD97) followed by the number of the marine isotopic stage (1 or 2)
and numbered from the bottom to the top (MD31-2-1 to MD31-2-5 and MD97-1-1 to MD97-1-6).
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F. Naughton, 2007
Pollen
zones
Pollen signature
MD97-1-6
Strong increase of Pinus (15-70%).
Continuous decrease of deciduous Quercus.
Ericaceae (55%), Poaceae (10%) and Taraxacum-type (10%).
MD97-1-5
Continuous decline of Pinus (40-15%),
deciduous Quercus (40-15%), Corylus and evergreen Quercus,
presence of Alnus (8-12%). Ericaceae increase (30-55%).
MD97-1-4
Gradual decline of Pinus (60-40%) and deciduous Quercus (6040%), maximum expansion of Corylus (6-10%),
beginning of Alnus continuous presence.
Gradual increase of Ericaceae (10-30%) and decrease of
herbaceous pollen percentages: Poaceae (<10%), Calluna
(<2%), Aster-type (<1%) and Cyperaceae (<2%).
Semi-desert (<3%). Spores presence (pislete triletes and Isoetes)
MD97-1-3
Pinus decline (~60%).
Maximum expansion of deciduous Quercus (60-80%), beginning
of Corylus continuous presence, presence of evergreen
Quercus.
Herbaceous pollen percentages decrease: Ericaceae (<10%),
Poaceae (<10%), Calluna (<2%), Aster-type (<1%) and
Cyperaceae (<2%), Semi-desert (<3%)
MD97-1-2
Pinus (80-90%).
Decrease of deciduous Quercus (40%) and increase of Betula
(~10%). Poaceae increase (20-30%), Ericaceae (10-20%),
Taraxacum-type (<10%). Increase of semi-desert associations:
Artemisia (~5-15%), Chenopodiaceae (~3%), Ephedra (~2%).
Younger
Dryas
(YD)
MD97-1-1
Pinus (80-90%).
Strong increase of tree percentages:
deciduous Quercus (40-60%).
Decrease of ubiquist associations: Poaceae (10-20%), Ericaceae
(<20%), Cyperaceae (~10%); Calluna (<5%). Presence of
pioneer species: Betula, Cupressaceae and Hippophae.
BöllingAllerød
(B-A)
MD31-2-5
Pinus (~ 80%).
Poaceae (30%), Ericaceae (10-15%), Calluna (<10%),
Cyperaceae (5-10%), Aster-type (~10%), Taraxacum-type
(~10%). Semi-desert associations: Artemisia (2-12%),
Chenopodiaceae (~3%), Ephedra (<2%).
Presence of pioneer species (Betula and Hippophae) (<5-10%).
MD31-2-4
Strong decrease of Pinus (~20-40%).
Poaceae (20-45%), Ericaceae (~20%), Calluna (10-20%),
Cyperaceae (<10%), Aster-type (10-20%),
Taraxacum-type (15-20%). Semi-desert associations: Artemisia
(<5%), Chenopodiaceae (<5%), Ephedra (~2%).
MD31-2-3
Pinus (~ 60%).
Poaceae (20-40%), Ericaceae (20-45%), Calluna (2-15%),
Cyperaceae (5-10%), Aster-type (2-15%),
Taraxacum-type (5-30%). Semi-desert associations:
Artemisia (<5%), Chenopodiaceae (<3%), Ephedra (<2%).
Presence of temperate trees (<5-10%) and pioneer species.
MD31-2-2
Pinus (30-40%).
Poaceae (20-40%), Ericaceae (20-30%),
Taraxacum-type (10-20%), Calluna (15-20%), Aster-type (10%).
Semi-desert associations:
Artemisia (<5%), Chenopodiaceae (1-2%).
MD31-2-1
Pinus (~ 60%).
Poaceae (0-20%), Ericaceae (20-30%), Calluna (10%),
Cyperaceae (5-10%), Aster-type (<10%),
Taraxacum-type (10-20%).
Semi-desert associations:
Artemisia (1-5%), Chenopodiaceae (0-3%), Ephedra (<2%).
Chronostratigraphy
Oldest
Dryas
Late Glacial period
Holocene
Late
Pleniglacial
Tab. II.3| Description of the pollen zones in the Galician margin composite core and respective
chronostratigraphy.
131
F. Naughton, 2007
2. 4. 3. 1. 1 Heinrich events (H2 and H1)
Our Galician margin composite record reveals two periods marked by
the dominance of herbaceous communities (Poaceae, Ericaceae, Calluna,
Cyperaceae, Aster-type, Taraxacum-type) along with a Pinus forest reduction
indicating two major cold events in north-western Iberia. These events, pollen
zones MD31-2-2 and MD31-2-4, are centred at around 21 700 yr BP and 14 700
yr BP, respectively. In the ocean, our record identifies H2 and H1 events on the
basis of, as usual in other North Atlantic cores, peaks in ice rafted detritus
(IRD), high polar foraminifera (N. pachyderma s.) percentages and heavy
planktonic δ18O values (e.g., Heinrich, 1988; Bond et al., 1993; Duplessy et al.,
1993; Grousset el al., 1993; Bond and Lotti, 1995; Lebreiro et al., 1996; Baas et
al., 1997; Abrantes et al., 1998; Cayre et al., 1999; Bard et al., 2000; Shackleton
et al., 2000; Thouveny et al., 2000; Broecker and Hemming, 2001; de Abreu et
al., 2003; Hemming, 2004). Radiocarbon ages obtained for H2 (~22 000 to ~20
000 yr BP) and H1 (~15 350 to ~13 000 yr BP) intervals in our Galician margin
record are in agreement with the age limits of these events, proposed by Elliot
et al. (1998) for the North Atlantic.
Direct correlation between pollen and marine proxies performed in this
record (Fig. II.6; Tab. II.3) shows that these major cold events in north-western
Iberia are only associated with the first part of H2 and H1. Indeed, H2 and H1
encompass two vegetational phases. Besides the Pinus forest contraction, the
first part of H2 (~22 000 to 21 500 yr BP; MD31-2-2 pollen zone) and H1 (~15 350
to 14 500 yr BP; MD31-2-4 pollen zone) is characterised by the expansion of
Calluna. Calluna vulgaris is a light demanding species (Calvo et al., 2002)
favoured by forest regression and moist conditions. This indicates that the first
part of both Heinrich events was cold and humid. Furthermore, the first part of
H1 is marked by the continuous presence of the Isoetes fern suggesting also
moist conditions.
The second part of H2 (21 500 to 20 000 yr BP; first 1 500 yr of the MD312-3 pollen zone) and that of H1 (14 500 to 13 000 yr BP, MD31-2-5) are marked
by a Pinus expansion, indicating less cold conditions than the previous
phases. Furthermore, during the second part of H1, a gradual increase of
semi-desert plants (Artemisia, Chenopodiaceae and Ephedra) reflects a
132
F. Naughton, 2007
gradual dryness on land. Our multiproxy palaeoclimatic record (Fig. II.6; Tab.
II.3) indicates therefore that H2 and H1 events display a complex pattern on
the adjacent continent.
This two-phase climatic succession on land within H2 and H1 agrees
with the changes detected in the marine proxy data from the same record
(Fig. II.6). The first phase is represented by the heaviest δ18O values of G.
bulloides and the increase of N. pachyderma (s.) percentages, suggesting a
strong decrease in sea surface temperatures (SST), and the absence of IRD in
this region. In contrast, the second phase records the lightening of the
planktonic isotopic signal and the decrease in the polar foraminifera
population although the presence of IRD testifies to iceberg melting off
Galicia. This complex marine pattern within H2 and H1 events has already
been detected further south in the SU81-18 deep-sea record (Fig. II.1) (Bard et
al., 2000).
Comparison between our multiproxy palaeoclimatic record (Fig. II.6)
and the available terrestrial and marine pollen sequences in and off Iberia
(Fig. II.1) indicate that the impact of H2 and H1 events in Iberia is spatially
variable. The Pinus reduction associated with H2, dated in the Galician
margin record between 22 040±180 and 21 340±160 years BP, has been
already detected by terrestrial and marine pollen sequences in and off
northern Iberia (Fig. II.1) (Pérez-Obiol and Julià, 1994; Roucoux et al., 2005). In
north-eastern Spain, the decrease in Pinus percentages before 19,900 yrs BP
detected in the Banyoles sequence and explained as the result of local
factors (Pérez-Obiol and Julià, 1994) can now be interpreted as the
consequence of the climatic change associated to H2. The slight expansion
of Calluna recorded in the Galician margin sequence during the first phase of
H2 has not been detected, however, at low altitude in north-eastern Spain
where Poaceae was the dominant taxa. This suggests that heathers grow
preferentially in the north-western Iberia favoured probably by Atlantic wet
conditions.
Southern Iberian margin cores (Fig. II.1) reveal the expansion of semidesert associations (Artemisia, Chenopodiaceae, Ephedra), suggesting an
increase of dryness during the entire H2, 22 000-20 000 years BP, although no
decrease of Pinus forest was detected (SU81-18, Turon et al., 2003; ODP 976,
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F. Naughton, 2007
Combourieu-Nebout et al., 2002 and SO75-6KL, Boessenkool et al., 2001). In
contrast, the Padul record (Pons and Reille, 1988; Fig. II.1; Tab. II.4) shows
between 23 600 and 19 800 yrs BP an alternation between periods of high
Pinus pollen values and periods of high percentages of semi-desert plants. This
suggests dryness variability in Sierra Nevada at that time or changes in polleninput related with local factors.
The two-phase climatic succession of H1 event characterised in our
Galician margin record (MD31-2-4 and MD31-2-5 pollen zones, Fig. II.6; Tab.
II.3) by a first cold and humid episode followed by a dry and cool phase is
contemporaneous, as H2, with a unique aridity interval and no Pinus forest
reduction in south-western Iberia (Fig. II.1) (Boessenkool et al., 2001; Turon et
al., 2003). Seemingly, high altitude sites of northern Iberia and eastern Iberian
sites (Padul and Banyoles) detect one event of dryness between 15 000 and
13 000 (Tab. II.4, Figs. II.1 and II.7; Laguna de la Roya, Allen et al., 1996;
Quintanar de la Sierra, Peñalba et al., 1997; Laguna Masegosa, Von
Engelbrechten, 1998; Lagoa Lucenza, Muñoz Sobrino et al., 2001; Laguna
Lleguna and Laguna de las Sanguijuelas; Muñoz Sobrino et al., 2004). The
westernmost sequences record the expansion of Calluna and Isoetes, as we
observed in the first part of H1 of our record, showing that these sites are also
affected by wet Atlantic influence (Lagoa Lucenza, Muñoz Sobrino et al.,
2001; Laguna Lleguna and Laguna de las Sanguijuelas, Muñoz Sobrino et al.,
2004; Mougás, Gómez-Orellana et al., 1998). Pinus forest cover around all high
altitude sites remains weak over this time-interval while our Galician margin
record (Figs. II.6 and II.7; Tabs. II.3 and II.4) representing also the vegetation
of low and mid altitudes sees a Pinus expansion in the second part of H1. This
suggests that the temperature increase was not enough to trigger Pinus
expansion in high altitude areas (Fig. II.1). These vegetational changes
related with cold conditions in Iberia coincide with the Oldest Dryas originally
identified in Danish deposits one century ago and dated older than 13 000
years BP (Mangerud et al., 1974) and not with the Older Dryas as erroneously
correlated by Turon et al. (2003). Therefore, our work demonstrates that the
Oldest Dryas is the terrestrial counterpart of the H1 event.
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F. Naughton, 2007
Years 14 C BP/
continental
sequences
Laguna de la Roya (1608 m a.s.l)
(Allen et al., 1996)
yr BP
Quintanar de la Sierra (1470 m a.s.l.)
(Penalba, 1994; Penalba et al., 1997)
yr BP
Padul (785 m a.s.l.)
(Pons and Reille, 1988)
yr BP
1200 -present
day
Local increase of Betula (30-50%).
Poaceae (20-30%), Ericaceae (5-10%),
Rumex (~2%).
Culture presence: Olea, Castanea and
Cerealia.
Absence of Pinus and Quercus decrease
Pinus, Fagus
and herbs (Ericaceae, Cerealia)
---
3000-1200
Poaceae (30-40%) well represented,
spread of Ericaceae (10%).
Decrease of Pinus, Betula and slight
decrease of Quercus
Spread of Fagus
---
3060
Spread of Corylus
Maximum percentages of trees
(Betula, Pinus, deciduous Quercus, Quercus
ilex and Corylus)
4450
8200
Succession of Juniperus, Betula and
deciduous Quercus and Q. ilex
8200
Deciduous Quercus, Quercus ilex, Pistacia
Pinus is almost absent (<5%)
10120
Pinus presence (20-40%)
Poaceae (20-30%), Artemisia (10%),
Apiaceae (10-15%), Plantago (5%), Aster
(2%), Cyperaceae (2-10%),
Chenopodiaceae and Calluna
10000
Younger Dryas
Poaceae (20-30%), Artemisia (10%),
Chenopodiaceae, Plantago,
Caryophyllaceae, Anthemis-type and
Calluna)
Decrease of trees until 40%.
Artemisia (>10%), Poaceae (20%),
Chenopodiaceae (>10%), Ephedra (5%),
Cyperaceae (20-30%)
Slight increase of Pinus
Late Glacial
interstadial
Succession of Juniperus, Betula, Quercus.
Pinus well represented (40-60%)
11050
Succession of Juniperus, Salix, Betula
Pinus well represented (60-80%)
Poaceae (20-40%) , Artemisia (20-40%),
Chenopodiaceae (2-5%) , Plantago (~2%),
Caryophyllaceae (2%), Aster-type (2%) and
Calluna (2%)
13350
6000-3000
Succession of Juniperus, Betula, Quercus,
Corylus and Alnus
10000-8200
10290
Late Glacial
(Q S and LR)/
Oldest Dryas
(Padul)
12940
Poaceae (20-40%), Artemisia (15-30%),
Chenopodiaceae (~5%), Plantago (0-7%),
Cyperaceae (5-10%), Aster (2-5%)
---
Juniperus, Betula, deciduous Quercus,
Quercus ilex and Pistacia
13200
Pinus decrease (<40%)
Artemisia (20%), Poaceae (20-40%),
Chenopodiaceae (>10%), Cyperaceae (60100%)
15200
Pinus increase (50-75%)
Artemisia (10-20%), Poaceae (10%),
Chenopodiaceae (<10%), Cyperaceae (030%). Presence of trees (5%)
19800
Alternation of coldest episodes: Artemisia
(30-60%), Chenopodiaceae (10%),
Cyperaceae (<20%), Poaceae (<10%) and
with less cold episodes: Artemisia (10-20%),
Chenopodiaceae (<5%), Cyperaceae (5080%), Poaceae (20%) and Pinus increase
(50-70%)
---
Late Pleniglacial
---
Deciduous Quercus, Quercus ilex, Quercus
suber, Pistacia.
Pinus is almost absent (<5%)
--23600
Tab. II.4| Description of pollen zones from the well-dated reference sites of Quintanar de la Sierra (Peñalba 1994, Peñalba et al., 1997), Laguna de la Roya (Allen et al., 1996) and
Padul (Pons and Reille, 1988).
135
F. Naughton, 2007
Fig. II.7 | - Comparison between continental (Quintanar de la Sierra; Peñalba et al., 1997) and marine
(MD99-2331 and MD03-2697) pollen sequences.
136
F. Naughton, 2007
2. 4. 3. 1. 2 The LGM
In our record, the Last Glacial Maximum (LGM), bracketed by H2 and
H1 events as established by the EPILOG program (Environmental processes of
Ice age: Land, Oceans, Glaciers) (Mix et al., 2001), is characterised by the
expansion of Pinus in an herbaceous-dominant environment along with
scattered pockets of deciduous trees (MD31-2-3 pollen zone) (Fig. II.6, Tab.
II.3). Planktonic δ18O values are slightly lower than during H2 and H1 events
and N. pachyderma s. decrease to very low percentages. Pinus percentages
stay constant over this interval and there is an almost continuous slight
presence of deciduous tree pollen (deciduous Quercus, Betula, Corylus and
Alnus) over this period. Nevertheless herbaceous communities remain the
dominant group. The presence of deciduous trees has also been detected in
and off southern Iberia between 20 000 and 15 000 yr BP (Tab. II.4; Pons and
Reille, 1988; Boessenkool et al., 2001; Combourieu-Nebout et al., 2002)
associated with the LGM (Turon et al., 2003). Our record clearly shows that not
only southern but also north-western Iberia acted as a refugium zone for
certain temperate trees (deciduous Quercus, Corylus, Alnus and Betula)
during the last glacial maximum corroborating what has been suggested by
previous studies (Roucoux et al., 2005). However, it must be noted that
deciduous trees presence is weak and that they attain their maximum
expression in southern Iberia. Another interesting feature within the LGM
concerns the sustaining of Ericaceae communities in north-western Iberia
indicated by our Galician margin core and their slight expansion in southern
Iberia detected by marine cores SU81-18 (Turon et al., 2003) and ODP 976
(Combourieu-Nebout et al., 2002). This is contemporaneous with the slight
decrease of semi-desert associations in the middle altitudes of Sierra Nevada
(Padul), indicating an increase of humidity in Iberia at that time.
2. 4. 3. 2. 1 The Bölling-Allerød
Following the H1 event, a drastic change in the pollen assemblage
and planktonic stable oxygen isotopic values identifies the Bölling-Allerød (BA) temperate period (Greenland Interstadial 1- GIS 1, Lateglacial interstadial).
137
F. Naughton, 2007
Our Galician margin pollen record (Fig. II.6, Tab. II.3) (MD97-1-1 pollen
zone) detects a fast deciduous Quercus expansion and the slight
development of pioneer species (Betula, Cupressaceae and Hippophae), a
decrease of herbaceous associations and Pinus percentages reach
maximum values. In the ocean, surface waters show an important lowering of
the δ18O values suggesting, in absence of freshwater input, an oceanic
warming at these North Atlantic mid-latitudes.
In the northern Iberian Península, the Lateglacial interstadial (B-A) is
characterised by the succession of pioneer associations (Juniperus-BetulaPinus) and the more or less important expansion of deciduous trees (Fig. II.1
and II.7; Tab. II.4; Allen et al., 1996; Peñalba et al., 1997; Von Engelbrechten,
1998). High altitude sites of northern Iberia such as Laguna Masegosa (Von
Engelbrechten, 1998), Laguna de la Roya (Allen et al., 1996), Hojos de Iregua
(Gil-Garcia et al., 2002) and Laguna de las Sanguijuelas (Muñoz Sobrino et al.,
2004) record a deciduous Quercus expansion between 13 000 and 11 000 yrs
BP above 1 000 m a.s.l.. Our pollen analysis records higher pollen percentages
of deciduous Quercus than the high altitude sequences suggesting that
deciduous Quercus woodlands expanded preferentially in the lowlands and
mid-altitudes of northern Iberia. Brewer et al. (2002), based on a small number
of high altitude northern and low altitude southern sequences, suggested that
only southern Iberia acted as a refugium zone for deciduous oak during the
last glacial period. However, as previously shown by our Galician record, the
mid and low-altitudes of north-western Iberia were a refugium zone for
deciduous Quercus species allowing the fast spread of these taxa during B-A
climate improvement.
In southern Iberia, a rapid expansion of deciduous and evergreen
Quercus and other Mediterranean elements is recorded in the Lateglacial
interstadial of the Padul peat-bog sequence and marine records SU81-18
(Turon et al., 2003), 8057 B (Hooghiemstra et al., 1992), SO75-6KL (Boessenkool
et al., 2001) and ODP 976 (Combourieu Nebout et al., 1999, 2002). Indeed, a
distinct phase of pioneer trees is not reflected at the beginning of this
interstadial neither in southern Iberia nor in low and mid altitudinal sites of the
north-western Iberia as shown by our Galician margin pollen record.
138
F. Naughton, 2007
2. 4. 3. 2. 2 The Younger Dryas cold event
Following the B-A warm phase, our Galician margin pollen record
(MD97-1-2 pollen zone) sees the increase of pioneer species (Betula), grasses
and semi-desert associations (Artemisia and Ephedra) at the expense of the
temperate forest. These vegetational features characterise the Younger Dryas
cold event (Fig. II.6, Tab. II.3).
This cold episode has been detected in and off Iberia (Pons and Reille, 1988;
Pérez-Obiol and Julià, 1994; Allen et al., 1996; Peñalba et al., 1997; von
Engelbrechten, 1998; Gil Garcia et al., 2002; Turon et al., 2003) (Fig. II.1, Tab.
II.4) and is associated with a slight increase of the planktonic δ18O values (Hall
and McCave, 2000; Schönfeld and Zahn, 2000; Löwemark et al., 2004; Turon
et al., 2003).
The slight decrease of deciduous Quercus, the increase of pioneer
species (Betula) and the expansion of semi-desert and herbaceous plants are
also represented in north-eastern Iberia as well as in continental and offshore
southern Iberian sequences (Pons and Reille, 1988; Pérez-Obiol and Julià,
1994; Boessenkool et al., 2001; Turon et al., 2003). Vegetation changes related
to this short event appear more drastic at high altitudinal sites of northern
Iberia than at low and mid altitude sites of north-western Iberia (this study)
and those from the south as indicated by the slight decrease of temperate
trees in terrestrial and marine pollen sequences in and off southern Iberia (Fig.
II.7).
2. 4. 3. 2. 3 The Holocene
After the YD, the tree succession of deciduous Quercus, Corylus and
Alnus defines the Holocene in our Galician margin record (Fig. II.6, Tab. II.3).
Remarkably, during the onset of the Holocene (MD97-1-3 pollen zone), the
increase of deciduous Quercus pollen percentages tightly parallels the
lightening of δ18O values (Fig. II.6). However, deciduous forest attains its
maximum expression slightly before sea surface water experiences its lightest
δ18O values. Pinus pollen values decline steadily in the pollen zones MD97-1-3
through MD97-1-5. The end of the maximum expansion of deciduous Quercus
trees and the beginning of the Corylus increase (MD97-1-4) occur
139
F. Naughton, 2007
contemporaneously with the beginning of the lightest δ18O isotopic values in
the ocean. Deciduous Quercus gradually decreases until the end of the
Holocene (from MD97-1-4 to MD97-1-6).
The expansion of heaths (Ericaceae and Calluna) and ferns (as
indicated by psilate Trilete spores) along with the reduction of trees
(deciduous Quercus, Corylus and Betula) mark the late Holocene phases
(MD97-1-5 and MD97-1-6). The minimum pollen values of Pinus are recorded in
zone MD97-1-5 while maxima pollen percentages of this tree, in the
successive zone MD97-1-6, probably reflect the reforestation of the last 350
years.
In
Iberia,
vegetation
response
to
climate
amelioration
that
characterises the Holocene period looks quite similar to that of the B-A event.
The settlement of the Mediterranean forest occurred very fast in southern sites
(Fig. II.1) as illustrated by the pollen sequences of Padul (Tab. II.4, Pons and
Reille, 1988), Charco da Candieira (Van der Knaap and Van Leeuwen, 1995),
SU81-18 (Turon et al., 2003), ODP 976 (Combourieu-Nebout et al. 2002), 8057 B
(Hooghiemstra et al., 1992), and SO75-6KL (Boessenkool et al., 2001). In the
north, vegetation response to the Holocene climate appears slower than in
the south. The expansion of pioneer trees (Juniperus, Betula and Pinus) marks
the beginning of this period in the high altitude sites of northern Iberia
followed by the development of deciduous Quercus, Corylus and Alnus (Tab.
II.5; Fig. II.1). This succession is also clearly detected by our Galician marine
record synchronously with the decrease of planktonic δ18O values which are
contemporaneous with the sea surface gradual warming detected by de
Abreu et al. (2003) and Schönfeld et al. (2003) in the Iberian margin. However,
this vegetation succession, recorded in our Galician sequence, begins earlier,
during the Younger Dryas event, in the low and mid altitude sites than in the
high altitudinal sites of north-western Iberia. The Galician margin record further
suggests that the maximum development of deciduous Quercus forest leads
the lightest values of planktonic δ18O during the Holocene. Finally, the late
Holocene interval of all Iberian marine and terrestrial sequences indicates the
decline of the temperate forest during the last 5000 years.
140
F. Naughton, 2007
Quintanar de
la Sierra
Laguna
de La
Roya
Lago
de Ajo
Laguna
Masegosa
Laguna Negra
Banyoles
Lake
Las Pardillas
Lake
(42°02’N,3°W)
(42°6’N,6°44’W)
(43°3’N,6°9’W)
(42°57’N,2°49’W)
(42°0’N,2°52’W)
(42°08’N,2°45’E)
(42°2’N,3°2’W)
1608 m asl
1085 m asl
1570 m asl
1600 m asl
1760 m asl
173 m asl
1850 m asl
Laguna
Lucenza
(Sierra de
Queixa)
(Galicia)
1420 m asl
Hoyos de
Iregua
Pozo do
Carballal
(42°01’N,2°45’W)
(42°42’N,7°07’W)
1780 m asl
1330 m asl
Lagoa Lucenza
(42°35’N,7°07’W)
1375 m asl
Laguna de las
Sanguijuelas
(42°08’N,6°42’W)
Holocene subphases
1080 m asl
Betula
Fagus
Late-Holocene
5000-4000-3000
yr BP
until present
day
Alnus
Ulmus
Corylus
Fraxinus
Alnus
Ulmus Corylus
Ulmus
Alnus
Fraxinus
Corylus
Mid-Holocene
9000-8000 yr BP
to
5000-4000 yr BP
Eve.Quercus
Dec.Quercus
Eve.Quercus
Dec.Quercus
Eve.Quercus
Dec.Quercus
Eve.Quercus
Dec.Quercus
Pinus
Betula
Salix
Betula
Juniperus
Pinus
Betula
Pinus
Juniperus
Salix
Juniperus
Pinus
Betula
Pinus
Betula
Betula
Pinus
Pinus
Pinus
Pinus
Betula
Pinus
Betula
Fagus
Taxus
Fagus
Taxus
Fagus
Fagus
Fagus
Taxus
Abies
Fagus
Salix
Fagus
Fagus
Corylus
Alnus
Ulmus
Fraxinus
Corylus
Ulmus
Alnus
Fraxinus
Corylus
Alnus
Ulmus
Taxus
Fraxinus
Corylus
Alnus
Ulmus
Fraxinus
Corylus
Alnus
Corylus
Alnus
Ulmus
Fraxinus
Corylus
Alnus
Ulmus
Corylus
Alnus
Salix Corylus
Eve.Quercus
Dec.Quercus
Dec.Quercus
Quercus
Eve.Quercus
Dec.Quercus
Eve.Quercus
Dec.Quercus
Eve.Quercus
Dec.Quercus
Eve.Quercus
Dec.Quercus
Dec.Quercus
Pinus
Salix
Betula
Juniperus
Betula
Juniperus
Betula
Juniperus
Pinus
Salix
Betula
Juniperus
Pinus
Salix
Betula
Juniperus
Acer
Betula
Pinus
Juniperus
Salix
Betula
Pinus
Juniperus
Salix
Betula
Pinus
Early Holocene
10500 yr BP to
9000-8000 yr BP
Tab. II.5| Holocene tree succession in north-western Iberia.
141
F. Naughton, 2007
2. 5 Conclusions
The comparison of present-day terrestrial and marine pollen samples in
and off western Iberia shows that the pollen signature from the Iberian margin
is similar to that of the Iberian terrestrial deposits, and, in particular, to that of
the estuarine samples which recruit pollen from the vegetation colonising the
adjacent hydrographic basins. Therefore, western Iberian margin pollen
spectra reflect an integrated image of the regional vegetation of the
adjacent continent. Furthermore, our study shows that marine pollen spectra
clearly discriminate both the Mediterranean and the Atlantic plant
communities colonising southern and northern Iberian Península, respectively.
It also identifies the present-day pattern of pollen transport in northern and
southern Iberian margin during downwelling and upwelling conditions.
High resolution pollen and marine proxies analysis from the Galician
margin composite core (MD99-2331 and MD03-2697) shows a synchronicity of
the vegetation response to the North Atlantic climatic variability during H2,
LGM, H1, B-A, YD events. Comparison of this palaeoclimatic record with other
marine and terrestrial pollen records shows that the beginning of both H2 and
H1 cold events are associated with Pinus forest reduction in northern Iberia. It
also shows the presence of two vegetation phases within H1 and H2 events,
associated with an initial cold and wet episode followed by a cool and,
particularly, dry episode during H1. Furthermore this comparison allows us to
demonstrate that the Oldest Dryas event on the continent corresponds to the
H1 event in the ocean.
The slight presence of deciduous Quercus, Corylus and Alnus during
the Last Glacial Maximum shows that not only southern Iberia but also
northern Iberia acted as a refugium zone for these trees, though at a smaller
scale. Bölling-Allerød interstadial in our sequence, which mainly represents low
and mid-altitude zones, show a more rapid and great expansion of
deciduous Quercus than the high altitude sites of north-western Iberia,
indicating that the vegetation of low and mid-altitudes responded more
rapidly to the climate variability of the North Atlantic during this interstadial.
Because deciduous forest attained it maximum expression during the B-A
interstadial in low and mid-altitudes of the north-western Iberia, the climate
reversal characterizing the Younger Dryas event is less marked in these zones
than in the high altitude ones. The response of deciduous forest to the climate
improvement that characterises the onset of the Holocene at low and mid
altitudes of north-western Iberia seems to lead those observed in the high
altitude sites, although the same succession of trees is observed in all these
northern regions.
This study confirms that marine pollen sequences from western Iberian
margin are a powerful tool for accurate reconstruction of vegetation
response to oceanic and atmospheric climate changes within a reliable
chronological framework. Furthermore, it demonstrates the importance of
including vegetation reconstructions from marine pollen sequences in future
efforts to refine and model vegetation and climate dynamics in the Iberian
Península.
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Capítulo 3| New insights on the impact of Heinrich events
and LGM in the mid-latitudes of the eastern North Atlantic
and in the adjacent continent
Novos conhecimentos sobre o impacto dos eventos de
Heinrich e do último máximo glaciar nas latitudes médias do
Atlântico Nordeste e no continente adjacente
Nouvelles approches sur l’impact des évènements d’Heinrich
et du Dernier Maximum Glaciaire dans les moyennes
latitudes dans l’est de l’Atlantique Nord et sur le continent
adjacent
A reduced version of this chapter will be submitted in December 2006 to:
Earth and Planetary Science Letters
F. Naughton a, b, M.F. Sánchez Goñi a, J. Duprat a, E. Cortijo c, B. Malaizé a, C. Joly a, S.
E. Bard d, F. Rostek d and J-L. Turon a
a
Environnements et Paléoenvironnements Océaniques (UMR CNRS 5805 EPOC), EPHE,
b
c
d
CEREGE, UMR-CNRS 6635, Aix-en-Provence, France
153
F. Naughton, 2007
Resumo
O estudo multidisciplinar de alta resolução da sondagem marinha,
MD99-2331, recolhida no noroeste da margem Ibérica, mostra que a resposta
da vegetação temperada e do oceano das médias latitudes do Atlântico
Norte ao último máximo glaciar (LGM) é contraditório ao que foi observado
durante os episódios interestadiais do MIS3 tardio (MIS3-Marine isotopic Stage
3). O aumento do albedo, o forte contraste sazonal das altas latitudes do
Atlântico Norte e a diminuição da concentarção de CO2 na atmosfera,
durante o LGM, podem ter sido os principais responsáveis pelo fraco
desenvolvimento da floresta temperada nesta região. Contudo, o aumento
da intensidade da circulação termohalina do Atlântico Norte (MOC)
favoreceu a transferência de humidade para o noroeste da Península
Ibérica provocando o desenvolvimento de Ericaceae e Calluna nesta
região.
Este estudo mostra ainda duas fases principais na vegetação Ibérica
as quais parecem estar intimamente ligadas ao complexo sinal deixado
pelos típicos eventos de Heinrich (H4, H2 and H1) ao longo desta margem. A
primeira fase, anterior à chegada máxima de icebergues a esta margem, é
marcada por condições oceânicas de superfície muito frias (evidenciada
pelas associações de foraminíferos, δ18O e estimativa da temperatura
baseada nas alcanonas) assim como por um arrefecimento extremo no
continente revelado pelo forte declínio da floresta de Pinus. O aumento de
Calluna e elevadas concentrações polínicas indicam o aumento das
condições húmidas durante esta primeira fase. A segunda fase, associada à
chegada máxima dos icebergs nesta margem, é caracterizada por
condições oceânicas de superfície e continentais menos frias e pelo
aumento da aridez a qual é representada pelo desenvolvimento de plantas
semi-desérticas.
O drástico arrefecimento detectado durante os eventos H4, H3, H2 e
H1 foi provavelmente provocado pela interrupção da MOC seguida de
rápidas reorganizações entre o oceano e a atmosfera as quais favoreceram
a transmissão das condições frias para o noroeste Ibérico. Para além deste
mecanismo
oceanográfico,
variações
semelhantes
ao
actual
modo
negativo e positivo do índice da Oscilação Norte Atlântica (NAO) parecem
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F. Naughton, 2007
ter tido um papel crucial no padrão climático deixado pelos eventos de
Heinrich no noroeste Ibérico. Durante a primeira fase, a predominância do
modo negativo da NAO-like gerou um aumento da precipitação no inverno
e importantes descargas fluviais nesta região. Estas condições favoreceram
ainda a fusão dos icebergues na cintura de IRD (IRD belt), onde a
temperatura da água superficial seria relativamente quente, impedindo o
seu transporte para as latitudes médias do Atlântico Norte. Durante a
segunda fase, a predominância do modo positivo da NAO-like provocou a
intensificação e a migração para norte dos ventos de oeste provocando um
aumento da aridez ao longo da Península Ibérica. Estas condições
favoreceram ainda a migração dos icebergues para sul e a sua fusão nas
latitudes médias do Atlântico Norte.
O evento atípico H3 reflecte condições húmidas durante quase todo
o episódio sugerindo uma não intensificação ventos de oeste nesta região.
Résumé
L’étude multiproxy et à haute résolution de la carotte marine profonde
MD99-2331 prélevée au nord-ouest de la marge Ibérique montre un
découplage entre la réponse de la forêt tempérée et les températures
relativement élevées des eaux océaniques de surface (SST) pendant le
Dernier Maximum Glaciaire (LGM) contrairement à ce qui est détecté
pendant les plus récents interstadiaires de Dansgaard-Oeschger du Stade
Isotopique Marin 3 (MIS3). L’augmentation de l’albédo, le fort contraste
saisonnier ainsi que la diminution de la concentration de CO2 pourraient
expliquer le faible développement des arbres tempérés dans cette région.
Cependant, l’augmentation de l’intensité de la circulation méridienne
Atlantique de renversement (MOC), transportant l’humidité vers la marge
nord-ouest Ibérique, a sûrement joué un rôle dans l’augmentation de SST et le
développement observé des bruyères.
Cette étude met également en évidence deux phases principales de
développement de la végétation dans le nord-ouest de la Péninsule Ibérique
intimement liées à la complexité de l’impact des évènements typiques
d’Heinrich (H4, H2 et H1) sur la marge Ibérique. La première phase, précédant
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F. Naughton, 2007
l’arrivée maximale d’icebergs sur cette marge, est marquée par des
températures des eaux de surface particulièrement froides, détectées par les
assemblages de foraminifères planctoniques, les analyses de δ18O et les
mesures de Uk37, associées au refroidissement important sur le continent
adjacent détecté par la réduction de la forêt de Pin. L’expansion de Calluna,
conjointement à l’augmentation de concentration pollinique, indique une
augmentation de l’humidité pendant cette phase. La deuxième phase,
associée à l’arrivée massive d’icebergs dans la marge Ibérique, est
caractérisée par des eaux de surface et des conditions atmosphériques
moins froides et par une augmentation de l’aridité identifiée par le
développement des plantes semi-désertiques.
Les conditions froides au cours des événements d’Heinrich 4, 3, 2 et 1
ont été probablement engendrées par l’arrêt de la MOC, suivi de
réorganisations rapides entre l’océan et l’atmosphère favorisant le transfert
instantané de conditions froides dans le nord-ouest de la Péninsule Ibérique.
Derrière ce mécanisme océanographique, des changements d’index
(négatif ou positif) de l’Oscillation Nord Atlantique (NAO-like) semblent avoir
joué un rôle crucial dans le scénario climatique complexe laissé par les
évènements d’Heinrich dans cette région. En effet, pendant la première
phase, un mode dominant d’index négatif de la NAO-like aurait généré,
comme c’est le cas à présent, une augmentation des précipitations
hivernales et donc des décharges des rivières dans la Péninsule Ibérique. Ces
conditions favorisent la fonte des icebergs au niveau de la ceinture de dépôt
préférentielle d’IRD où les eaux de surface sont relativement chaudes,
empêchant leur migration vers le sud jusqu’aux moyennes latitudes. Pendant
la seconde phase, une plus forte fréquence des situations de NAO-like en
mode positif aurait conduit à une intensification des vents d’ouest et leur
déplacement vers le nord, produisant une augmentation de la sécheresse
dans la Péninsule Ibérique. Ces conditions auraient favorisé la migration vers
le sud des icebergs vers les moyennes latitudes.
Par ailleurs, l’évènement atypique H3 est caractérisé dans cette région
par des conditions humides pendant presque toute sa totalité, certainement
liées au maintien sur cette région de vent d’ouest affaiblis.
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Abstract
High resolution multi-proxy study of MD99-2331 deep sea core retrieved
in north-western Iberian margin shows that the response of temperate forest
and mid-latitude sea surface temperatures to the Last Glacial Maximum
(LGM) period was decoupled contrarily to what has been detected during
the late MIS 3 interstadials. Albedo increase, high seasonality and CO2
concentration decrease could explain the weak development of temperate
trees in this region. However, the more vigorous Meridional overturning
circulation (MOC) transferring moisture to north-western Iberia could be
responsible for the observed heathland development.
Furthermore, this study evidences two main vegetation phases in northwestern Iberia linked to the complex imprint left by the typical Heinrich events
(H4, H2 and H1) in the Iberian margin. The first phase, before the maximal
arrival of icebergs into this margin, is marked by extremely cold sea surface
temperatures indicated by the planktonic foraminifera assemblages and δ18O
analyses together with Uk37 measurements and by the strong cooling of the
adjacent continent revealed by the Pinus forest decline. Calluna expansion in
concert with the highest total pollen concentration indicates moisture
increase during this phase. The second phase, associated with the maximum
arrival of icebergs into the Iberian margin, is characterised by less cold sea
surface and atmospheric conditions and by an increase of dryness identified
by the development of semi-desert plants.
The cold conditions during Heinrich events 4, 3, 2 and 1 were probably
triggered by the Atlantic MOC shutdown followed by ocean-atmosphere
rapid reorganizations favouring the transfer of cold conditions into northwestern Iberia. Besides this oceanographic mechanism, changes similar to
that of prevailing (negative and positive) North Atlantic Oscillation (NAO)
index seems to have played a crucial role on the complex climatic pattern
left by Heinrich events in north-western Iberia. Indeed during the first phase,
prevailing negative mode of NAO-like index generates winter precipitations
and river flow increase in Iberia. These prevailing conditions, favoured iceberg
melting in the IRD belt (where sea surface temperature was relatively warm)
preventing their southern migration to the mid-latitudes. During the second
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phase, prevailing positive mode of NAO-like index leads to westerlies
intensification and northward displacement triggering an increase of dryness
in Iberia. These prevailing conditions favoured the southward migration of the
icebergs to the mid-latitudes sites.
The atypical H3 is characterised in this region by wet conditions over
almost the entire event probably due to maintaining reduced westerlies in this
region.
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3. 1 Introduction
In the last years, many studies have been carried out to understand
the sources, trigger mechanisms and the global impact of the well known
massive episodes of iceberg discharges into the North Atlantic that occurred
during the last glacial period (see Hemming, 2004). These extreme episodes,
named Heinrich events, have been firstly documented in the Ruddiman belt
from several North Atlantic deep-sea cores between 45 and 50° N (Heinrich,
1988, Bond and Lotti, 1995). They were identified by the anomalous presence
of ice-rafted detritus (IRD) that were transported to the ocean by drifting
icebergs from Laurentide and northern European ice sheets (Heinrich, 1988)
as well as by peaks of N. pachyderma (s) (e.g. Bond and Lotti, 1995;
Hemming, 2004) and magnetic susceptibility (Grousset et al., 1993). These
coarse fraction intervals, representing the well known IRD layers, were also
detected out of the Ruddiman belt i.e. north of 50°N (e.g. Fronval et al., 1995;
Rasmussen et al., 1996; Elliot et al., 1998; Voelker et al., 1998; Van Kreveld et
al., 2000) as well as below 40°N (e.g. Lebreiro et al., 1996; Baas et al., 1997;
Zahn et al., 1997; Chapman et al., 2000; Bard et al., 2000; de Abreu et al.,
2003). The thickness of the IRD layers and the magnetic signal is, however,
smaller in the mid-latitude sites than in the northern ones (Thouveny et al.,
2000; Dowdeswell et al., 1995). Also, the duration of the impact of these
extreme events on the sea surface temperatures (SST) is in this region longer
than that of the IRD layers (e.g. Bard et al., 2000; Chapman et al., 2000;
Sánchez Goñi et al., 2000). Indeed, the Heinrich events have left a complex
pattern imprint along the Iberian margin (e.g Abrantes et al., 1998; Bard et al.,
2000; Thouveny et al., 2000; Schönfeld et al., 2003; Narciso et al., 2006).
Western Iberian vegetation further shows a complex pattern
associated with Heinrich events (H) 2 and 1 (Naughton et al., 2006) and with
H3, H4 and H5 (Sánchez Goñi et al., 2000). The first hypothesis based on landsea direct correlation has linked the relatively wet conditions in south western
Iberia at the beginning of H5 to H3 events to “European iceberg discharges”,
and the semi-desert plant development to the successive massive iceberg
discharges from the Laurentide ice sheet (Sánchez Goñi et al., 2000). This
extreme dryness has been attributed to a prevailing positive North Atlantic
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Oscillation (NAO) index (Sánchez Goñi et al., 2002). However, the
mechanisms proposed for explaining this complex pattern within Heinrich
events in the Iberian margin are not conclusive so far.
Most of the direct correlations between terrestrial and marine proxies
from Iberian margin deep-sea cores does not deeply discuss about the Last
Glacial maximum (LGM) period. However, the LGM is considered as a key
interval for understanding the sensitivity of global environmental systems to
change (Mix et al., 2001), because it represents the extreme opposite
situation to an interglacial and a period of relatively stable glacial maximum
conditions. Recently, several sea surface temperature (SST) reconstructions for
the LGM period have been carried out around the world by MARGO project
(Multiproxy Approach for the Reconstruction of the Glacial Ocean surface)
suggesting the presence of seasonal ice cover in the North Atlantic and in the
Nordic Seas (de Vernal et al., 2005). This sea-ice free season would allow the
Meridional Overturning Circulation to supply moisture to the northern
hemisphere high latitudes (Meland et al., 2005). Previous SST reconstructions
carried out in the Iberian margin have clearly demonstrated that during the
LGM sea surface conditions were warmer than during Heinrich events (e.g.
Bard et al., 1987; Lebreiro et al., 1997; Cayre et al., 1999; Bard et al., 2000;
Pailler and Bard, 2002; de Abreu et al., 2003). Because LGM SST values are
similar or even higher than those characterising Late Marine Isotopic Stage 3
(MIS3) Dansgaard-Oeschger (D-O) interstadials, we should expect similar
amplitude of temperate trees expansion in north-western Iberia what it is not
the case. Indeed, north-western Iberian margin MD95-2039 record (Roucoux
et al., 2005) does not reveal the same amplitude of temperate tree expansion
during both the MIS 3 D-O interstadials and the LGM. Therefore, other
mechanisms than the weakening of the MOC has to be seek to explain the
weak development of temperate trees in this region during the LGM.
The aim of this work is firstly to propose the possible trigger mechanisms
for the complex pattern of vegetation observed in north western Iberia during
Heinrich events. Secondly, we will discuss the paradox observed during the
LGM when relatively high SST in the mid-latitudes of the eastern North Atlantic
were associated with cool environments on land.
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MD99-2331 deep sea core was recovered on the Galician margin (42°
09’ 00 N, 09° 40’ 90 W) at ~100 km from the coast and at 2110 m water depth
(Fig. III.1). Morphology, recent sedimentation and hydrology of the Iberian
margin have been thoroughly described in Naughton et al. (2006).
The northernmost part of north-western Iberia is influenced by the wet,
cool and weakly seasonal Atlantic climate (annual precipitation = 900-1400
mm and annual temperature ranges = -7 and 10° C). This region is dominated
by deciduous Quercus forest (Q. robur, Q. pyrenaica and Q. petraea), heaths
communities (Ericaceae and Calluna) and Ulex (Alcara Ariza et al., 1987;
Polunin and Walters, 1985).
Fig. III.1 | Map showing MD99-2331 location and sites of the cores referred in the text: 1: MD95-2040 (Pailler
and Bard, 2002; de Abreu et al., 2003; Schönfeld et al., 2003; Narciso et al., 2006), 2: MD95-2039 (Thouveny
et al., 2000; Roucoux et al., 2001; 2005; Schönfeld et al., 2003), 3: PO 28-1 (Abrantes et al., 1998), 4: D11957P
(Lebreiro et al., 1996; 1997), 5: SO75-26KL (Zahn et al., 1997; Boessenkool et al., 2001), 6: PO 8-2 (Abrantes et
al., 1998), 7: MD95-2042 (Cayre et al., 1999; Sánchez Goñi et al., 2000; 2002; Thouveny et al., 2000; Pailler
and Bard, 2002), 8: SU81-18 (Bard et al., 2000; Turon et al., 2003); 9: ODP 976 (Combourieu et al., 2002), 10:
SU90-03 (Chapman et al., 2000), 11: ESSCAMP-KS02 (Loncaric et al., 1998; Zaragosi et al., 2001), 12: MD952002 (Grousset et al., 2000; Zaragosi et al., 2001), 13: AKS01 (Zaragosi et al., 2001), 14: VM 23-81 (Bond and
Lotti, 1995), 15: MD04-2845 (work in progress), 16: SU90-11 (Jullien et al., in press.), 17: MD03-2705 (Jullien et
al., submitted), 18: OCE326-GGC5 (McManus et al., 2004).
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F. Naughton, 2007
Core MD99-2331 was retrieved using a giant CALYPSO corer during the
Ginna (IMAGES V) oceanographic cruise on board the R/V Marion Dufresne
(Fig. III.1). This sedimentary record, mainly composed of hemi-pelagic clay, is
37.2 m long and covers the last Marine Isotopic Stages (MIS) 7 to 1. In this
study, we will focus on the last 40 000 years where sedimentary rates vary from
43 cm kyr-1 to 30 cm kyr-1, providing a high-resolution palaeoclimatic record
for this time period off north-western Iberia and in the adjacent continent.
X-ray analysis on this core using SCOPIX image-processing (Migeon et
al., 1999) shows a rather well preserved sedimentary sequence between 2
and 11 m core depth.
3. 3. 1 Chronostratigraphy
The age model of MD99-2331 deep sea core is based on 55
accelerator mass spectrometer (AMS)
14C
dates obtained at “Laboratoire de
Mesure du Carbone 14” (LMC) in Saclay and at AMS laboratory (GifA) in Gifsur-Yvette on monospecific samples with maxima of Globigerina bulloides or
Neogloboquadrina pachyderma (s.) abundances (Fig. III.2 and Tab. III.1).
AMS
14C
younger than 21 786 BP were calibrated using CALIB Rev 5.0
program and the "global" marine calibration dataset (marine 04.14c) (Stuiver
and Reimer, 1993; Hughen et al., 2004; Stuiver et al., 2005). We used the 95.4%
(2 sigma) confidence intervals and their relative areas under the probability
curve as well as the median probability of the probability distribution (Telford
et al., 2004) as suggested by Stuiver et al. (2005).
14C
radiometric ages older than 21 786 yr BP and younger than 40 000
yr BP were firstly corrected for the regional marine age reservoir of -400 yr and
then, calibrated using a simple second-order polynomial (Age cal yr BP = 6.2724 x 10-6 x [Age
14C
yr BP]2 + 1.3818 x [Age
14C
yr BP] – 1388) which has
been constructed by means of Iberian margin data tuned with GISP2 (Bard et
al., 2004).
All conventional radiocarbon dates as well as their calibration results
are plotted against core depth in Fig. III.2. We use both the paleoclimate
162
F. Naughton, 2007
multi-proxy record of MD99-2331 as well as the age limits of Heinrich events
proposed by Elliot et al., 2002 to delimit H4, H3, H2 and H1 events.
Some levels associated with Heinrich events reflect age inversions
(older and younger) when compared with adjacent host sediments (Fig. III.2
and Tab. III.1). They were rejected because they likely reveal reworking
processes
due
to
substantial
changes
in
sedimentary
rate.
Others,
representing bioturbated levels (presence of zoophycos burrows) with
younger material as well as those which are too old compared with Elliot et
al. (2002) age limits were also rejected. Finally, we have also rejected some
anomalous dates associated with isolated peaks of IRD that occurred after H4
and H3.
Therefore we only use 29 AMS
14C
levels for establishing the chronology
of MD99-2331 during MIS2 and late MIS3 (Fig. III.2 and Tab. III.1).
Fig. III.2 | Chronostratigraphy of the MD99-2331 record. AMS radiocarbon dates are represented by
triangles while the calibrated ones are represented by squares. North Atlantic Heinrich events are delimited
by both the age limits (not calibrated) proposed by Elliot et al. (2002) and by the results obtained from the
multi-proxy study of the MD99-2331 record (see below). White triangles and squares reflect the rejected
levels for the model age while the dark ones represent the accepted ages.
163
F. Naughton, 2007
Lab code
Core
depth
(cm)
Material
Conv.
AMS
14C
age BP
Conv.
AMS 14C
age BP
(-400 yr)
error
95.4 % (2σ)
Cal BP
age ranges
Cal BP
age
Median
probability
LMC14-001231
LMC14-001232
GIF-102377
LMC14-001233
LMC14-001235
LMC14-001236
LMC14-001237
LMC14-002445
GIF-101109
GIF-102373
LMC14-002446
LMC14-001845
LMC14-001846
LMC14-001847
LMC14-001849
LMC14-001850
GIF-102378
LMC14-001851
LMC14-001852
LMC14-001853
LMC14-001854
LMC14-001855
GifA-102374
LMC14-001856
LMC14-002447
LMC14-001857
LMC14-001858
LMC14-002448
LMC14-001859
LMC14-001860
LMC14-001861
GifA-102375
LMC14-001862
LMC14-001863
LMC14-001864
LMC14-001865
LMC14-001866
LMC14-001867
LMC14-001868
LMC14-001869
LMC14-001870
LMC14-001871
LMC14-001872
GifA-102376
LMC14-001873
LMC14-002449
LMC14-001874
LMC14-001875
LMC14-001876
LMC14-001877
LMC14-001878
GifA-102379
LMC14-001879
LMC14-001880
LMC14-002450
200
205
220
222
228
235
242
260
290
570
590
595
600
607
620
623
630
637
650
655
670
700
740
740
760
770
780
800
810
820
830
840
850
860
870
880
890
895
920
925
945
960
970
980
985
990
995
1000
1005
1010
1015
1020
1025
1030
1040
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
G. bulloides
G. bulloides
G. bulloides
G. bulloides
G. bulloides
G. bulloides
G. bulloides
G. bulloides
G. bulloides
N. pachyderma
G. bulloides
G. bulloides
G. bulloides
G. bulloides
G. bulloides
G. bulloides
G. bulloides
N. pachyderma
G. bulloides
G. bulloides
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
G. bulloides
G. bulloides
G. bulloides
G. bulloides
N. pachyderma
N. pachyderma
G. bulloides
G. bulloides
G. bulloides
N. pachyderma
G. bulloides
G. bulloides
G. bulloides
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
N. pachyderma
G. bulloides
13640
13810
14130
13920
13930
15130
15060
15540
16170
19770
22290
20860
20550
20460
21620
21740
22690
22150
22440
22430
23080
24140
25880
25140
26310
23960
24390
25870
26050
26370
27010
27790
28530
28510
28860
29030
29470
29540
30260
30020
31960
32720
32530
33410
34440
34810
34490
35350
34600
31490
32950
33850
34920
36940
37100
13240
13410
13730
13520
13530
14730
14660
15140
15770
19370
21890
20460
20150
20060
21220
21340
22290
21750
22040
22030
22680
23740
25480
24740
25910
23560
23990
25470
25650
25970
26610
27390
28130
28110
28460
28630
29070
29140
29860
29620
31560
32320
32130
33010
34040
34410
34090
34950
34200
31090
32550
33450
34520
36540
36700
80
80
120
90
80
90
90
90
130
170
170
250
240
140
160
160
180
170
170
180
190
210
250
230
260
200
210
250
280
260
280
320
350
350
370
380
390
390
430
420
560
570
570
500
700
700
700
790
710
510
620
440
760
920
940
15303:16099
15524:16359
15898:16828
15658:16520
15686:16521
17250:18182
17170:18038
18405:18723
18787:19265
22534:23622
15679
15922
16342
16067
16081
17848
17722
18520
18983
23038
23931:25369
23450:24803
23652:24405
25301:26000
25439:26000
24542
24119
24016
25626
25730
Bard
et al.,
2004
25854
26296
25699
26020
26009
26725
27881
29748
28959
30204
27686
28151
29737
29928
30267
30940
31754
32519
32498
32858
33032
33480
33552
34280
34038
35974
36720
36534
37390
38381
38733
38428
39244
38533
35509
36944
37815
38837
40728
40876
Tab. III.1| Radiocarbon ages of MD99-2331 deep sea cores. Bold levels represent the accepted ages while
not bold ones represent the rejected ages for the age model.
164
F. Naughton, 2007
MD99-2331 core was sub-sampled for pollen analysis between 2.00 and
11.00 m. The sample spacing was between 2 and 5 cm for MIS2 and 5 to 30
cm for the late MIS3, allowing a time resolution average of 110 years and 210
years, respectively.
The Galician margin is influenced by north-western prevailing winds
which impede substantial direct airborne transport of pollen grains from the
continent to the sea. Because, north-western Iberia is composed of a
complex fluvial system, pollen grains are mainly released by the Douro river
followed by the Minho river and transferred seaward following a estuary-shelfslope transect (Naughton et al., 2006). Furthermore, the comparison between
modern terrestrial and marine pollen signals in and off western Iberia
demonstrate that marine pollen spectra reflect an integrated image of the
regional vegetation of the adjacent continent (Naughton et al., 2006).
Sample preparation technique follows de Vernal et al. (1996) modified
at the UMR CNRS 5805 EPOC (Desprat, 2005). An exotic, Lycopodium, spike of
known concentration has been added to each sample to calculate total
pollen concentrations (including Pinus, aquatic plants, spores, indeterminate
and unknown pollen grains).
After chemical treatments (cold 10%, 25% and 50% HCl as well as cold
40% and 70% HF) the samples were sieved through a 10 µm nylon mesh
screen (Heusser and Stock, 1984) and mounted in bidistillate glycerine. Pollen
and spores were counted using a Zeiss Axioscope light microscope at x550
and x1250 (oil immersion) magnifications. A minimum of 100 pollen grains
excluding the over-represented Pinus grains in the Iberian margin (Naughton
et al., 2006) have been counted. At least 20 taxa and 100 Lycopodium were
also counted in each of the 165 samples analysed. Pollen percentages were
calculated based on the main pollen sum which excludes Pinus, aquatic
plants, spores, indeterminate and unknown pollen grains.
Detailed description of pollen data from MIS2 has been previously
published in Naughton et al. (2006).
165
F. Naughton, 2007
3. 3. 3 Marine proxy analysis
A total of 101 oxygen isotopic measurements were carried out on
Globigerina bulloides planktonic foraminifera (Gouzy et al., 2004; Naughton et
al., 2006) at Bordeaux University with a sample spacing of 2 to 20 cm (time
resolution average of 240 years). Sixty-one measurements were made on
Cibicides wuellerstorfi benthic foraminifera every 2 to 90 cm (time resolution
average of 400 years) at the Laboratoire des Sciences du Climat et de
l’Environnement (LSCE), Gif-sur-Yvette, France.
Each specimen has been picked up within the 250–315 µm fraction
and cleaned with distilled water. The preparation of each aliquot (4–10
specimens with 80 μg of weight) has been carried out in the Micromass
Multiprep autosampler by using an individual acid attack. The CO2 gas
extracted has been analysed against NBS 19 standard, taken as an
international reference standard. Planktonic Isotopic analysis has been
carried out using an Optima Micromass mass spectrometer at the UMR CNRS
5805 EPOC and benthic isotopic analysis were performed using a delta plus
Finnigan at the LSCE. The mean external reproducibility of powdered
carbonate standards is ±0.05‰ for oxygen. Results from oxygen isotopic
analysis are presented versus PDB.
3. 3. 3. 2 Ice rafted detritus (IRD)
188 levels (2 to 10 cm sample spacing with an age resolution average
of 130 yr) were washed with distilled water and wet sieved trough a 150 μm
mesh
screen.
After
this
classic
sedimentological
procedure,
IRD
semiquantitative analysis was performed on the >150 μm sand-size fraction. In
this study, only the total concentrations of the lithic grains were considered.
The presence of two anomalous small peaks of IRD after H4 and H3
associated with age inversions likely shows reworked levels as the result of
destabilizations of the slope through changes of sea-level (Figs. III.2 and III.4).
3. 3. 3. 3 Planktonic foraminifer-derived SST
154 levels (2 to 20 cm of sample spacing with a time resolution average
of 150 years) were treated for planktonic foraminifera analysis. Sub samples
166
F. Naughton, 2007
were washed with distilled water and wet sieved trough a 150 mesh screen.
At least 400 specimens per sample were counted for semiquantitative analysis
and identified based on Kennet and Srinivasan (1983).
Planktonic foraminifera were grouped into three main bioclimatic
assemblages: polar (Neogloboquadrina pachyderma sinistral), subpolar
(Neogloboquadrina
pachyderma
dextra,
Globigerina
bulloides
and
Turborotalia quinqueloba) and warm including temperate/cold, subtropical,
warm subtropical and tropical, (Globorotalia scitula, G.inflata, G.hirsuta,
G.truncatulinoides,
G.crassaformis,
Globigerinita
glutinata,
Globigerina
falconensis, G.calida,G.rubescens, G.digitata, Hastigerina aequilateralis,
Orbulina universa, Globigerinoides ruber) (e.g. Bé, 1977; Ottens, 1991; Duprat,
1983).
Winter (February) and summer (August) sea-surface temperatures SST
were estimated by using the modern analogue technique transfer function
from the database of Pflaumann et al. (1996) improved by E. Cortijo (LSCE,
Gif-sur-Yvette, France) and J. Duprat (UMR CNRS 5805 EPOC) on planktonic
foraminifera assemblages.
3. 3. 3. 4 Alkenone-derived SST
Long-chain C37-C39 ketones, or alkenones are biomarkers synthesised by
algae (Prymnesiophyceae class) such as the coccolithophores E. huxleyi and
Gephyrocapsa oceanica (Volkman et al., 1980; Volkman et al., 1995). The
unsaturation ratio of C37 alkenone (Uk37΄) is directly correlated with the
temperature of Emiliania huxleyi growth in laboratory cultures and therefore, it
is used as proxy for sea surface temperature estimations (Prahl et al., 1988;
Rostek et al., 1993; Rosell-Melé et al., 1995).
Calibration on Uk37΄ had been applied to modern sediments (Sikes et
al., 1991; Rosell-Melé et al., 1995) agreeing with previous calibration on the E.
huxleyi cultures. We will assume that our data represent the annual SST even if
the impact of seasonality on the SST signal remains unsolved (Sachs et al.,
2000).
136 levels (5 to 10 cm of sample spacing) were used for alkenone
analysis providing a time resolution average of 180 yr. The extraction of
alkenone from the sediment was performed using an automated Dionex
167
F. Naughton, 2007
Accelerated Solvent Extractor (ASE-200) in Cerege, Aix-Marseille III, CNRS
UMR-6635. C37 concentration was determined using n-C36 added to the ASEcell before the extraction processes and SST values were calculated based
on Prahl et al. (1988) equation.
The weak C37 concentration detected at around 10 m core depth
prevented reliable SST estimates for H4 event.
3. 4 Results and discussion
Direct correlation of both continental and marine high-resolution
records from the MD99-2331 deep-sea core allows us to recognise the
vegetation changes in north-western Iberia linked with Heinrich events 4 to 1
and the LGM period, and to discriminate the main mechanisms underlying
these changes.
3. 4. 1 Long-term climate variability and the LGM period
Both the late MIS3 (starting at around 40 000 cal yr BP) and almost the
entire MIS2, representing the final stages of the last glacial period, are
characterised by the gradual increase of the global ice extent which is
indirectly testified by the steady increase of heavy benthic δ18O values in the
MD99-2331 record (Fig. III.3) as previously suggested by Shackleton (1987).
This gradual increase of the global ice is synchronous with the long-term
Greenland atmosphere temperature decrease (Sánchez Goñi et al., in prep.)
during a period characterised by weak high-latitude summer insolation
(Berger, 1978). The gradual long-term cooling is synchronous with the longterm pattern of Pinus and temperate forest contraction as well as Poaceae
expansion in north-western Iberia (Fig. III.3). The general trend of gradual
forest tree decline has been previously detected in this region (Roucoux et al.,
2001; Roucoux et al., 2005) and elsewhere in Europe further south (Sánchez
Goñi et al., 2000) and east (Tzedakis et al., 2004).
168
F. Naughton, 2007
Fig. III.3 | Comparison between long term trends of MD99-2331 record and Greenland temperatures
(Sánchez Goñi et al., in prep.) during the Late MIS 3 and MIS 2 against age (cal yr BP). From the bottom to
the top: MD99-2331 benthic δ13C; percentages of Pinus, temperate and humid trees and Poaceae and
Greenland temperatures. Dashed line represents the long term trend of each signature.
The maximum ice sheet extension took place between 30 000 yr and
19 000 yr contemporaneously with a minimum of the sea-level (Lambeck et
169
F. Naughton, 2007
al., 2002). This period is however punctuated by a number of D-O and H
events. Its steady state is only observed in the interval bracketed between H2
and H1 events (Mix et al., 2001). This period of relatively stable climate which
was dramatically different from that of today, representing maximum glacial
conditions, has been defined by EPILOG as the LGM.
Paradoxically,
the
LGM
interval
in
the
MD99-2331
record
is
characterised by a huge decrease of polar foraminifera and an increase of
sub-polar and warm planktonic assemblages suggesting an increase of SST
values in the north-western Iberian margin at this time (Fig. III.4). Planktonic
foraminifera climate reconstruction estimate 14° to 17°C and 9° to 13°C for
summer and winter SST, respectively, while alkenone-based estimates show
annual SST means of about 12°-13°C (Fig. III.4).
Although LGM SST values are similar or even higher than those
characterising the late MIS3 D-O interstadials (GIS8 to GIS3) in the MD99-2331
record, north-western Iberia vegetation did not reply in the same way (Figs.
III.3 and III.4). During GIS8, 7, 6 and 5, temperate and humid trees slightly
developed in this region but more than during GIS 4 and 3 (Fig. III.4). They
were slightly present over the LGM (Fig. III.4). It has been proposed (Sánchez
Goñi et al., 2000; Sánchez Goñi, 2006) that during D-O interstadials of MIS3
temperate forest expansion started when summer SST reached 12°C. At
present, the distribution of temperate forest in both sides of the North Atlantic
coincides with summer SST between 12 and 18°C (Van Campo, 1984).
Because throughout the LGM period summer SST off north-western
Iberian was higher than 12°C we should expect the expansion of temperate
forest in this region. However, the vegetation cover of north-western Iberia
during the LGM period was dominated by pine and herbaceous communities
(including heaths and central-European steppe species) and by the slight
presence of temperate and humid trees, suggesting that only scattered
pockets of deciduous trees colonised this region (Naugthon et al., 2006).
Furthermore, here and in south-western Iberia (Turon et al., 2003) wet
conditions prevailed during the LGM likely as a response to a more vigorous
Meridional
Overturning
Circulation
(MOC)
testified
by
231Pa/230Th
measurements which estimate only less than 30-40% of MOC slowdown
170
F. Naughton, 2007
through this period (McManus et al., 2004) and predicted by numerical
climate models (Ganopolsky et al., 1998).
In the framework of the MARGO program, a compilation of several
climate reconstructions, including the Iberian margin region, further shows a
gradual latitudinal SST decrease from a southern to northern transect (e.g
Morey et al., 2005) during the LGM period. This latitudinal climatic gradient is
right expressed in pollen diagrams from this region (Boessenkool et al., 2001;
Turon et al., 2003; Combourieu-Nebout et al., 2002; Pons and Reille, 1988;
Roucoux et al., 2005; Naughton et al., 2006) suggesting that deciduous
Quercus forest expanded more easily in southern than in northern Iberia. Even
if this latitudinal SST gradient has contributed for the different amplitude
response between southern and northern vegetation during the LGM, the
magnitude of the temperate tree expansion is not linearly related to the SST
increase as it was during the D-O interstadials (Fig. III.3). Therefore other
mechanisms seem to have played a crucial role for precluding the large
expansion of temperate trees during the LGM in both regions.
One of the mechanisms that could explain the reduction of the
temperate tree forest, during this period, would be the increase of albedo
due to the maximum expansion of the northern ice-sheets and North Atlantic
sea-ice cover (Broccoli, 2000) which affected the steady-state global-mean
temperature (Rahmstorf, 2002). Indeed, the long term decline of temperate
and humid trees in western Iberia and the gradual amplitude decrease of the
D-O interstadials tightly parallels benthic foraminifera δ18O curve which
despite its low resolution clearly shows the general trend of ice volume
increase (Fig. III.3). This suggests that the ice accumulation via the albedo
feedback probably masked the impact of the mid-latitudes high SST values in
western Iberia during the LGM period. Moreover, this period is marked by the
increase of seasonal contrast as the result of sea-ice expansion during winter
and its reduction during summer in the North Atlantic and Nordic Seas (de
Vernal et al., 2005) which would also contribute to prevent deciduous trees
expansion in western Iberia. Also, the CO2 concentration average of 200
ppm during the LGM, representing 35% less than the present-day values
(Cowling and Sykes, 1999), could had play an important role for precluding
the expansion of temperate forest in western Iberia.
171
F. Naughton, 2007
Fig. III.4 | Multi-proxy results of MD99-2331 record. From bottom to top: ice-rafted detritus (IRD)
concentrations, percentages of planktonic foraminifera associations (polar, sub-polar and warm),
planktonic foraminifera-based winter and summer SST estimates, alkenone-based annual SST
reconstruction, δ18O of G. bulloides, percentages of temperate and humid trees, Pinus percentages and
Greenland temperatures (Sánchez Goñi et al., in prep.). Grey lines represent the Heinrich events. Note: the
age limits of H4 that are based on GISP2 chronology are slightly different from those estimated from the
calibration using NGRIP.
172
F. Naughton, 2007
3. 4. 2 Heinrich events
Superimposed to the late MIS3 and MIS2 long term-cooling, Heinrich
events, in particular H4, H3, H2 and H1, display a complex pattern in the northwestern Iberian margin (Fig. III.4), similar to that previously revealed by other
mid-latitudes North Atlantic deep sea cores (e.g. Zahn, 1997; Abrantes et al.,
1998; Loncaric et al., 1998; Bard et al., 2000; Chapman et al., 2000; Grousset
et al., 2000; Zaragosi et al., 2001; Thouveny et al., 2000; Schönfeld et al., 2003;
Narciso et al., 2006) (Fig. III.1).
Although IRD are not detected in the MD99-2331 record at the
beginning of the typical Heinrich events, a first drop of annual, summer and
winter SST estimates, the increase of planktonic δ18O values and that of the
planktonic polar foraminifera clearly show that these events have left an
imprint on north-western Iberian margin before the maximal arrival of icebergs
in this region (Fig. III.4). The evidence of this impact is reliable dated in northwestern Iberian margin, and are in agreement with the age limits proposed
by Elliot et al. (2002) elsewhere in the North Atlantic region.
North-western Iberia vegetation has reacted contemporaneously with
this complex pattern left by the Heinrich events in the Iberian margin. The
beginning of H4, H3, H2 and H1, marked by a SST drop in the ocean, is
synchronous with an important episode of Pinus forest contraction on land
suggesting a strong decrease of atmospheric temperatures (Fig. III.4). Northwestern Iberia atmospheric cold conditions are also testified by a drop in
temperate and humid trees even though they played a secondary role in the
total tree cover during the glacial period (Fig. III.4).
Pinus forest contraction occurred, therefore, slightly before the
maximal presence of IRD in the north-western Iberian margin during every
Heinrich events recorded in the MD99-2331 deep-sea core (Figs. III.4 and
III.5). Within H4 interval, the second Pinus forest contraction is, however,
synchronous with the maxima of IRD. Higher resolution analysis is required to
better constrain this episode. The multiproxy diagram of MD95-2039 deep sea
core (40° 34’ N, 10° 20’ W) (Roucoux et al., 2005) also illustrates that during
Heinrich events 5 to 1 the maximum of Pinus forest contraction occurred
before the maxima of IRD in this region. Pinus forest decline has been
173
F. Naughton, 2007
detected further north in MD04-2845 deep sea core retrieved in the Bay of
Biscay (work in progress).
Another important feature within the H4, H2 and H1 in north-western
Iberia is that the first phase is marked by an increase of Calluna and of total
pollen concentration, indicating wet conditions (Fig. III.5). Indeed, the sharp
decline of Pinus forest favoured continental erosion while the increase of
Calluna indicates an increase of moisture, intensifying river discharges and
seaward pollen transfer (Fig. III.5). Following this, the second phase is
characterised by the expansion of semi-desert plants reflecting an increase of
dryness contemporaneously with the maxima of IRD arrival in the northwestern Iberian margin and less cold sea surface conditions (Figs. III.4 and
III.5). The impact of H3 in north-western Iberia is peculiar and associated with
wet conditions over almost the entire event.
Fig. III.5 | Response of the north-western Iberia vegetation to the complex pattern of Heinrich events in the
Iberian margin. From bottom to top: Heinrich events, percentages of Pinus, percentages of Calluna
representing wet conditions, percentages of semi-desert plants reflecting continental dryness and total of
pollen concentration. Dashed line separated wet from dryness conditions during H events.
174
F. Naughton, 2007
Evidence for a gradual increase of dryness during each of the last five
Heinrich events have been detected further south in Iberia (Sánchez Goñi et
al., 2000) and in the Mauritanian margin MD03-2705 deep sea core (Jullien et
al., submitted) (Fig. III.1). Peaks in dust from this tropical core occur at the end
of the putative Heinrich events (Jullien et al., submitted). This suggests as our
record does that the mid- and low- latitudes of the eastern North Atlantic
region have experienced a first relatively wet period followed by increasing
dryness conditions during at least the last five Heinrich events. However, dust
content was lower during H3 than through H4, H2 and H1, suggesting that this
episode was less dry than the others also in the north-eastern tropical region.
3. 4. 3 Possible mechanisms triggering the complex pattern signal of
Heinrich events in and off north-western Iberia
Different hypotheses had been proposed to explain the complex
pattern observed during Heinrich events in the mid-latitudes of the North
Atlantic: a) multiple IRD sources (Bard et al., 2000; Thouveny et al., 2000;
Sánchez Goñi et al., 2000); b) multiple pulses of the Laurentide Ice sheet
(Abrantes et al., 1998), and c) shifts in the polar front position (Chapman et al,
2000). However, none of these hypotheses can explain either the succession
of wet and dry conditions during the typical H4, H2 and H1 or the drastic
atmospheric and sea surface cooling characterising the first stage of these
extreme events in north-western Iberia margin and adjacent landmasses.
Furthermore, the increase of dryness in Iberia coinciding with the maximum
dust input off Mauritania revealed by core MD03-2705 (Jullien et al.,
submitted) suggests that mechanisms triggering both atmospheric signatures
in temperate and tropical North Atlantic latitudes were probably connected.
Indeed, both regions are influenced by changes of the North Atlantic
Meridional Overturning Circulation (MOC) coupled with changes in prevailing
(negative or positive) mode of the North Atlantic Oscillation (NAO) index in
north-western Iberia and with shifts of the Intertropical Convergence Zone
(ITCZ) in north of Africa (Marshall et al., 2001).
Present-day relatively warm and moistness conditions of western
Europe are sustained by the poleward surface branch of the Atlantic Ocean
thermohaline circulation (THC) transferring heat from the tropics to the high
175
F. Naughton, 2007
latitudes of the northern hemisphere (Rahmstorf, 1995). However, this system is
highly sensitive to freshwater input which disturbs the strength of the THC
affecting regional and global climate (e.g., Rahmstorf, 1995, Broecker and
Hemming, 2001; Clark et al., 2002; Vellinga and Wood, 2002; Timmermann et
al., 2005).
In order to understand the impact of a Heinrich event on climate,
several climate models have introduced anomalous freshwater pulses into
the North Atlantic to force the MOC to shutdown (e.g. Seidov and Maslin,
1999, Ganopolski and Rahmstorf, 2001; Knuti et al., 2004; Rahmstorf et al.,
2005) triggering a substantial SST drop in that region (e.g. Paillard and
Labeyrie, 1994; Seidov and Maslin, 1999). Besides modelling simulations, the
episode of drastic and complete MOC shutdown has been confirmed for H1
by
231Pa/230Th
measurements from OCE326-GGC5 deep sea core (McManus
et al., 2004) (Fig. III.1). Following this, rapid oceanic and atmospheric
reorganizations favoured the transfer of atmospheric cold conditions to the
north-western Iberia triggering the Pinus and temperate forests decline in this
region.
MOC shutdown should also preclude moisture transfer to Europe during
the episode of Pinus forest decline. On the contrary, an increase of wet
conditions in north-western Iberia is observed. Therefore other mechanisms
must be invoked to explain the increase of wet conditions during the first
phase of Heinrich events in north-western Iberia. One of the mechanisms that
could explain the increase of wet conditions during the first phase of typical
Heinrich events is an atmospheric one related with weak pressure gradient
between high and low latitudes of the North Atlantic (negative mode of the
North Atlantic Oscillation-NAO). The north-western Iberian dryness during the
second part of typical Heinrich events would be, in turn, interpreted as a
prevailing positive mode of the NAO (Figs. III.6 and III.7).
Despite the importance of the NAO conditions upon present-day
wintertime climate it is still unknown how the NAO have changed in the past.
During the last glacial period the great extension of ice in the Northern
Hemisphere has probably affected the latitudinal and longitudinal position of
the low- and high-pressure cells. Therefore, the use of NAO mode changes as
an interpretive tool for explaining the complex climatic pattern of north-
176
F. Naughton, 2007
western Iberian Heinrich events must be made with caution. For this reason
we will use the term NAO-like from now on.
Indeed, during the wet first phase of a Heinrich event in this region both
the Azores high and the Icelandic low were weak, giving rise to reduced
westerlies over the eastern North Atlantic triggering prevailing negative NAOlike index conditions over Europe (Fig. III.6). This would favour, as at presentday (Trigo et al., 2004), the increase of winter precipitation and river flow in
western Iberia as revealed by the development of Calluna and the increase
of total pollen concentrations in our record. Furthermore, this atmospheric
mechanism produces nowadays warm sea surface conditions in the northwestern Atlantic above 45°N (Wanner et al., 2001) (Fig. III.6). During the first
phase of a Heinrich event, this warming might have facilitated iceberg
melting in the IRD belt as supported by the highest IRD concentrations (e.g.
Heinrich, 1988; Andrews and Tedesco, 1992; Grousset el al., 1993; Bond and
Lotti, 1995; Gwiazda et al., 1996; Hemming et al., 1998) and consequent
cooling of SST in that region. This likely hindered iceberg southern migration to
the mid-latitudes and explain the almost absence of IRD at the beginning of
H4, H3, H2 and H1 in western Iberia margin (Fig. III.6). The proposed prevailing
negative mode of the NAO-like index also favours the decrease of SST along
north-western Iberian margin until the Greater North Sea (Fig. III.6).
Fig. III.6 | Prevailing NAO negative conditions scheme (adapted from Wanner et al., 2001).
177
F. Naughton, 2007
The dry second phase of H4, H2 and H1 in Iberia, demonstrated by the
expansion of semi-desert plants from our MD99-2331 record was probably the
result of the strong pressure gradient between Açores and Iceland which
favoured the northward displacement and intensification of the westerlies
(Fig. III.7). At present-day, during prevailing positive mode of NAO, warm sea
surface conditions move towards the western mid-latitudes of the North
Atlantic (Wanner et al., 2001) (Fig. III.7). During the second phase of Heinrich
event the observed increase of IRD in the Iberian margin deep sea cores
could be explained by the warmth of mid-latitudes SST which favoured the
iceberg maximum arrival and melting at these latitudes (Fig. III.7). At this
moment most of deep sea cores from the IRD belt displays a gradual
decrease of IRD content. The prevailing positive mode of NAO-like index can
probably explain such a relatively warming of SST although iceberg melted at
this time in western Iberian margin.
Fig. III.7 | Prevailing NAO positive conditions scheme (adapted from Wanner et al., 2001).
To confirm that changes in the NAO modes operated within Heinrich
events we need to know whether there is an asymmetry between the east
and west of the North Atlantic climate during these events. Grimm et al (2006)
suggest based on Lake Tulane pollen record that Heinrich events in Florida
178
F. Naughton, 2007
were characterised by warm and wet conditions. However, the low resolution
data of Lake Tulane record and the radiometric dating-based correlation
between land and ocean climatic proxies precludes the identification of two
possible vegetation phases in this pollen sequence during Heinrich events.
Therefore, future work on pollen-rich deep-sea cores from the western North
Atlantic is crucial to support our hypothesis.
The impact of H3 in Iberia is different from H4, H2 and H1. Indeed H3
has been considered as an “atypical” Heinrich event (Snoeckx et al., 1999)
being characterised by a European ice sheet dominant signal (Bond et al.,
1992; Grousset et al., 1993; Snoeckx et al., 1999). Furthermore, studies carried
out in marine sediments from the high-latitudes of the North Atlantic and in
the western North Atlantic suggest that during H3 Laurentide ice sheet (Elliot
et al., 1998; Jullien et al., in press) as well as the others pan-Atlantic ice sheets
(Jullien et al., in press) may have discharged icebergs on a much smaller
scale than during the typical Heinrich events. The relatively wet conditions
detected in north-western Iberia (this study) and off Mauritanian coast (Jullien
et al., submitted) over the whole H3 would be the result of reduced wind-field
intensification in both regions likely due to a reduced pressure gradient
between Açores and Iceland in the North Atlantic region and to relatively
small southern migration of the ITCZ in the tropical zone.
3. 5 Conclusions
Direct correlation between marine and terrestrial proxies from MD992331 deep sea core shows that during the LGM temperate tree expansion
was largely reduced when compared with the previous late MIS 3 D-O
interstadials although SST were similar. We propose three mechanisms to
explain this decoupling between the ocean and atmosphere temperatures:
a) albedo increase which enhanced the cooling produced by low summer
insolation in the Northern Hemisphere, b) high seasonality, and c) weak CO2
concentration. Nonetheless, the wet conditions in Iberia during the LGM were
likely the result of the strengthening of the MOC which was more vigorous
than during the bracketing Heinrich events.
179
F. Naughton, 2007
During Heinrich events 4, 3, 2 and 1 the introduction of large amounts
of freshwater via Northern hemisphere icebergs drifting and consequent
melting triggered a shutdown of the Atlantic Meridional Overturning
Circulation (MOC) and a drop of SST in the North Atlantic region. This
produced ocean-atmosphere rapid reorganizations which allow the fast cold
conditions transfer into north-western Iberia triggering Pinus forest decline.
Superimposed to this important cooling, changes of prevailing
(negative and positive) North Atlantic Oscillation (NAO-like) index seems to
have played a crucial role for explaining the complex pattern of Heinrich
events in north-western Iberian margin and in the adjacent continent.
Indeed, Heinrich events in this region is characterised by two main
phases: a) the first is marked by a drop of SST and the virtual absence of
icebergs in the Iberian margin, an important fall of atmospheric temperatures
(strong Pinus forest contraction) and an increase of moisture conditions
(Calluna
expansion
in
concert
with
the
increase
of
total
pollen
concentration), b) the second phase is characterised by less cold conditions,
maximal arrival of icebergs and an increase of dryness (semi-desert plants
expansion). During the first phase of H4, H2 and H1, prevailing negative NAOlike index likely triggered the increase of winter precipitation in Iberia and
enhanced river flow favouring the seaward pollen transfer. Furthermore, these
prevailing conditions allowed iceberg melting in the IRD belt preventing their
southern migration to the mid-latitudes. This prevailing NAO-like mode also
favoured the drop of SST in north-western Iberian margin. Through the second
phase of H4, H2 and H1, prevailing positive NAO-like index conditions
intensified and moved northward the westerlies triggering an increase of
dryness in this region. The displacement of warm sea surface conditions to the
mid-latitudes of the North Atlantic favoured the southward migration of the
icebergs to the mid-latitudes sites. This mechanism also contributes to a
relatively warming of sea surface conditions in north-western Iberia. H3 is an
exceptional case of prevailing wet conditions during almost the entire event
probably due to maintaining reduced westerlies in this region.
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F. Naughton, 2007
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Capítulo 4| Climate variability during the last deglaciation
in north-western Iberian margin and adjacent continent
Variabilidade climática durante a última deglaciação no
noroeste da Península Ibérica e margem continental
adjacente
Variabilité climatique au cours de la dernière déglaciation
dans le nord-ouest de la marge Ibérique et sur le continent
adjacent
A reduced version of this chapter will be soon submitted to:
Paleogeography, Paleoclimatology, Paleoecology
F. Naughton a, b, M.F. Sánchez Goñi a, J. Duprat a, E. Cortijo c, S. Zaragosi
a
a
Environnements et Paléoenvironnements Océaniques (UMR CNRS 5805 EPOC), EPHE,
b
c
189
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Resumo
A correlação directa entre indicadores paleoclimáticos terrestres e
marinhos, obtidos a partir do estudo de uma sondagem marinha profunda
(MD03-2697), recolhida no noroeste da margem Ibérica, permitiu detectar
uma variabilidade climática sub-orbital, nas médias latitudes do Atlântico
Norte, durante a última deglaciação.
Durante o último máximo glaciar (LGM), o noroeste da Península
Ibérica foi afectado por um clima relativamente frio e húmido enquanto que
o oceano adjacente apresentava temperaturas de superfície relativamente
quentes. O aumento da intensidade da circulação termohalina do Atlântico
Norte (MOC) favoreceu a transferência de humidade para o noroeste da
Península Ibérica enquanto que o aumento do albedo, o forte contrate
sazonal das altas latitudes do Atlântico Norte e a diminuição do conteúdo
em CO2 atmosférico contribuíram para a manutenção de condições frias
sobre o continente.
Nas médias latitudes do Atlântico Norte, o evento de Heinrich 1 está
marcado por um padrão complexo, constítuido por duas fases distintas. A
primeira fase é marcada por condições oceânicas extremamente frias e
uma fraca presença de icebergues, enquanto que a segunda fase é
caracterizada por temperaturas da massa de água superficial menos frias e
por uma grande quantidade de detritos provenientes da fusão de
icebergues. No noroeste da Península Ibérica a vegetação respondeu
contemporaneamente a estas duas fases que caracterizam o evento de
Heinrich 1. O declínio da floresta de Pinus e a expansão de Ericaceae
(incluindo
Calluna),
durante
a
primeira
fase,
marcam
condições
extremamente frias e húmidas no noroeste da Península Ibérica, enquanto
que na segunda fase a expansão de plantas semi-desérticas e da floresta de
Pinus reflectem condições relativamente frias e áridas no continente. O
drástico arrefecimento que caracteriza a primeira fase foi provavelmente
provocado pela interrupção da MOC seguido por reorganizações rápidas
entre o oceano e a atmosfera, enquanto que as fases húmida e seca
resultam do predomínio de condições semelhantes aos modos negativo e
positivo do índice da oscilação Norte Atlântica (NAO-like), respectivamente.
190
F. Naughton, 2007
O aquecimento continental (expansão de Quercus deciduous) e
oceânico que caracteriza o evento Bölling-Alleröd (B-A) foi favorecido pelo
aumento da insolação de verão nas latitudes médias do Hemisfério Norte
assim como pela intensificação da MOC. Há cerca de 14 000 anos a
expansão máxima de Quercus deciduous a qual reflecte um aquecimento
continental extremo é síncrona do pico máximo na temperatura da
Gronelândia e do episódio de subida súbita do nível do mar (MWP 1A). Este
aquecimento severo do Hemisfério Norte poderá ter sido o impulsionador
deste drástico evento.
Durante o Dryas recente (YD), a diminuição da floresta de Quercus e a
expansão de plantas semi-desérticas reflectem um arrefecimento e aumento
da aridez continental contemporâneo da diminuição da temperatura da
massa de água superficial no oceano adjacente. A redução da intensidade
da MOC, embora associada ao máximo de insolação de verão das latitudes
médias do Hemisfério Norte, favoreceu uma diminuição em vez de um
declínio total da floresta decídua de Quercus no noroeste Ibérico. Para além
da redução da intensidade da MOC, a predominância de condições
semelhantes ao actual índice positivo da NAO (NAO-like) parece justificar o
aumento da aridez observado ao longo deste período.
Nesta região, o máximo térmico do Holocénico (HTM) foi detectado
entre 11 700 e 8 200 anos cal BP. Por volta dos 8 200 anos cal BP um rápido
arrefecimento continental (diminuição da floresta decídua de Quercus e
Corylus) e oceânico marca o evento frio “8.2 ky” no noroeste da Península
Ibérica e margem adjacente. Este episódio resulta do culminar de uma
sucessão de episódios relaccionados com o colapso da calote glaciar da
“Laurentide”
a
qual
provocou
a
amplificação
da
diminuição
da
temperatura na Europa e na Gronelândia.
Após o evento “8.2 ky” a diminuição gradual da floresta temperada é
contemporânea da diminuição da temperatura induzida pela diminuição
da insolação de verão das latitudes médias do Hemisfério Norte. Isto sugere
que a regressão da floresta temperada parece ter sido mais afectada pelas
variações orbitais do que pelo impacto antrópico.
191
F. Naughton, 2007
Résumé
La corrélation directe entre des indicateurs climatiques terrestres
(pollen) et marins d’une carotte marine profonde, MD03-2697, prélevée dans
le nord-ouest de la marge Ibérique a permis de détecter une variabilité
climatique d’ordre millénaire au cours de la dernière déglaciation dans les
moyennes latitudes de l’Atlantique Nord.
Le dernier maximum glaciaire (LGM) a été relativement froid et
humide dans le nord-ouest de la Péninsule Ibérique, alors que les
températures des eaux de surface étaient chaudes. L’intensification de la
circulation méridienne Atlantique de renversement (MOC) a favorisé le
transfert d’humidité des moyennes latitudes de l’Atlantique Nord vers la
marge ouest Ibérique tandis que l’augmentation de l’albedo, le fort
contraste saisonnier et la chute de la concentration de CO2 atmosphérique
ont maintenu des températures froides sur le continent.
L’événement d’Heinrich 1 (H1) a été marqué par un scenario
complexe dans les moyennes latitudes de l’Atlantique Nord. La première
phase est caractérisée par des températures des eaux de surface
extrêmement froides et par la quasi-absence d’icebergs alors que la
deuxième phase est légèrement moins froide et caractérisée par une forte
quantité de débris provenant de la fonte d’icebergs (IRD). La végétation du
nord-ouest de la Péninsule Ibérique a répondu de façon synchrone au
scenario complexe qui caractérise l’H1. Au cours de la première phase, la
réduction drastique de la forêt de pins et l’expansion des bruyères révèlent
des conditions extraordinairement froides et humides. La deuxième phase est
caractérisée par l’expansion de la forêt de pins et des plantes semidésertiques indiquant des conditions relativement froides et arides. Le
refroidissement qui chartérise la première phase a été probablement
engendrée par une coupure de la MOC, suivi par des réorganisations rapides
entre l’océan et l’atmosphère tandis que les phases humides et arides
résulteraient de situation dominante caractérisée par un index négatif et
positif de l’oscillation Nord Atlantique (NAO-like), respectivement.
Le
réchauffement
atmosphérique
et
océanique,
marqué
par
l’expansion de la forêt de chêne caducifolié et l’augmentation des
192
F. Naughton, 2007
températures des eaux de surface, caractérisant l’événement du BöllingAlleröd (B-A), est la conséquence de l’augmentation de l’insolation d’été des
latitudes moyennes de l’Hémisphère Nord et de l’intensification de la MOC.
L’expansion maximale de la forêt de chêne reflète un incident extrêmement
chaud vers 14 000 cal ans BP lequel est synchrone à la fois du réchauffement
maximal au Groenland et de l’épisode nommé Meltwater Pulse 1A (MWP 1A).
Ce réchauffement sévère de l’Hémisphère Nord pourrait être le principal
responsable de la fonte drastique et soudaine des calottes glaciaires.
La réduction de la forêt de chêne et l’expansion des plantes semidésertiques
pendant
l’événement
du
Dryas
récent,
indiquent
un
refroidissement et une augmentation de l’aridité sur le continent. Cet
événement est contemporain de la diminution des températures des eaux
de surface. La réduction de l’intensité de la MOC et l’augmentation de
l’insolation d’été des moyennes latitudes de l’Hémisphère Nord ont favorisé la
réduction plutôt que la disparition complète de la chênaie dans le nordouest de la Péninsule Ibérique. Au-delà de la diminution de l’intensité de la
MOC, la prévalence de l’index positif de la NAO-like pourrait expliquer
l’aridité détectée dans cette région.
Le maximum thermique de l’Holocène (HTM) à été détecté dans la
Péninsule Ibérique entre 11 700 et 8 200 cal ans BP. Un soudain refroidissement
continental (diminution de la forêt de chêne et du noisetier) et océanique
marque l’événement « 8.2 kyrs » dans cette région. Ce refroidissement serait
le résultat d’épisodes successifs associés à la réduction de la calotte
Laurentidienne produisant un fort refroidissement sur l’Europe et au
Groenland.
La diminution graduelle à long terme de la forêt tempérée après
l’événement « 8.2 kyrs » serait la réponse à un refroidissement induit par des
changements orbitaux plutôt qu’à un impact anthropique.
Abstract
Direct correlation between terrestrial (pollen) and marine climatic
indicators from deep sea core MD03-2697 (north-western Iberian margin)
193
F. Naughton, 2007
allows the detection of millennial scale climate variability for the last
deglaciation in the mid-latitudes of the North Atlantic realm.
The Last Glacial Maximum (LGM) was relatively cold and humid in
north-western Iberia while sea surface conditions were warm. More vigorous
Meridional Overturning Circulation (MOC) favoured moisture transfer from the
mid-latitudes of the North Atlantic to the western Iberia whereas increasing
albedo, high seasonality and atmospheric CO2 drop maintain the continent
cold.
The mid-latitudes of the North Atlantic were marked by a complex
pattern within Heinrich 1 (H1) event. In the first phase, sea surface conditions
were extremely cold with almost no evidence for iceberg calving while the
second one was less cold with high quantity of Ice-rafted detritus (IRD). In
north-western Iberia vegetation has responded synchronously to this H1
pattern. During the first phase a drastic Pinus forest decline and heaths
expansion reflect extremely cold and moist conditions whereas the second
phase reveals the expansion of Pinus forest and semi-desert plants
representing relatively cold and dry conditions on the continent. The first
coldest phase was probably triggered by the MOC shutdown followed by
ocean-atmosphere rapid reorganizations while the wet and dry phases were
the result of prevailing negative and positive North Atlantic Oscillation (NAOlike) indexes, respectively.
The continental (deciduous Quercus expansion) and sea-surface
warming characterizing the Bölling-Alleröd (B-A) event was produced by both
the increase of mid-latitude summer insolation of the northern Hemisphere
and the strengthening of the MOC. Maxima of deciduous Quercus expansion
reflecting extremely warming at 14 000 cal yr BP is synchronous with both the
peak of Greenland temperatures and the Meltwater Pulse 1A (MWP 1A). This
severe warmth of the Northern Hemisphere could be the trigger of this drastic
meting episode.
During the Younger Dryas (YD) the decrease of deciduous Quercus
forest and the expansion of semi-desert plants reflect continental cooling and
dryness which is contemporaneous with sea surface cooling. MOC reduction
but increasing northern mid-latitudes summer insolation favoured a decrease
rather than a complete decline of deciduous Quercus forest in north-western
194
F. Naughton, 2007
Iberia. Beyond the MOC reduction, a prevailing positive NAO-like index could
explain the observed dryness. Following this, the Holocene Thermal Maximum
in this region is identified between 11 700 and 8 200 cal yr BP. At around 8 200
cal yr BP a sudden land (decrease of deciduous Quercus forest and Corylus
woodlands) and sea cooling marks the 8.2 Ky event in Iberia as the result of
the culmination of the successive episodes of the Laurentide Ice sheet decay
which enhanced the cooling over Greenland and Europe. After the 8.2 Ky
event the long-term temperate forest decrease has responded to the
orbitally-induced cooling rather than to human impact.
195
F. Naughton, 2007
196
F. Naughton, 2007
4. 1 Introduction
Abrupt widespread millennial-scale climate changes during the last
deglaciation have been widely documented for the North Atlantic high and
mid-latitudes (e.g. Bond et al., 1993; Keigwin and Lehman, 1994; Andrews et
al., 1995; 1999; McManus et al., 2004) as well as from the tropics (Hughen et
al., 1996; Arz et al., 1999; Rühlemann et al., 1999; Peterson et al., 2000) and
Southern Atlantic Region (Kim et al., 2002; Shemesh et al., 2002). However, the
mechanisms
responsible
for
inter-hemispheric
and
tropical-pole
teleconnections are far from being completely understood (Hughen et al.,
1996; Blunier and Brook, 2001; Wunsch, 2003). Changes in heat transfer via
termohaline circulation, resulting from iceberg melting, are one of the
mechanisms proposed for explaining sub-orbital scale climate variability
(Bond and Lotti, 1995; Knuti et al., 2004; McManus et al., 2004) while others,
propose that these oscillations are a consequence of changes in the
topography of continental ice sheets leading to windfield shifts (Wunsch,
2006).
Among the mechanisms that appear to be involved in this sub orbital
climatic variability, shifts in the mode of prevailing North Atlantic Oscillation
(NAO-like) index coupled with changes in the Meridional Overturning
Circulation (MOC) could explain the complex climatic pattern of Heinrich
events during the late Marine Isotopic Stage (MIS) 3 and MIS 2 in northwestern Iberian margin and the adjacent continent (Naughton et al., in
prep.). However, these events occurred in a period characterised by a
relatively high and stable ice volume while the climatic variability of the last
deglaciation is affected by ongoing increase in summer insolation and
substantial reduction of ice volume in northern high latitudes. Previous works
have shown, for example, that the decrease in the MOC strength during the
Younger Dryas together with the insolation maximum lead to a mitigated cold
event (McManus et al., 2004).
To further explore the links between MOC strength and atmospheric
variability in the North Atlantic region superimposed to the gradual evolution
of insolation, we correlate north-western Iberia climatic and vegetation
197
F. Naughton, 2007
changes with North Atlantic mid-latitudes sea surface conditions and with
temperature estimates of Greenland.
Deep sea core MD03-2697 (42° 09’ 59 N, 59° 42’ 10 W) was retrieved at
~100 km off the Galician margin (north-west of Iberia) and at 2164 m of water
depth (Fig. IV.1). This site is at present day under the influence of the North
Atlantic Deep water mass. Morphology, recent sedimentation, detailed
hydrology and regional climate of this region have been previously described
in Naughton et al. (2006).
Briefly, north-western Iberia is influenced by a temperate and humid
climate (mean annual temperature of about 12. 5 °C and precipitation
varying from 1000 to 2000 mm) (Atlas Nacional de España, 1992); is incised by
two important seaward sediment and pollen suppliers such as the Douro river
followed by Minho, especially during downwelling conditions (Dias et al.,
2002; Jouanneau et al., 2002; Oliveira et al., 2002; Naughton et al., 2006) and;
is dominated by oak woodlands (Quercus robur, Q. pyrenaica and Q.
petraea), heaths communities (Ericaceae including Calluna), brooms
(Genista) and gorses (Ulex) (Alcara Ariza et al., 1987).
Fig. IV.1 | Study area. Location of deep-sea cores referred in the text: OCE326-GGC5 (McManus et al.,
2004); MD95-2002 (Zaragosi et al., 2001; Auffret et al., 2002; Ménot et al., 2006) and MD99-2331 (Naughton
et al., 2006; in prep.).
198
F. Naughton, 2007
Core MD03-2697 was recovered using a CALYPSO corer during PICABIA
oceanographic cruise on board the R/V Marion Dufresne (Fig. IV.1). MD032697, mainly composed of hemipelagic clays, is 41.23 m long covering the
last 425 000 years (from Marine Isotopic Stages (MIS) 1 to 11). In this study, we
will focus on the last deglaciation period.
Mean sedimentary rate is high (30 cm kyr-1) between ~ 20 000 yr cal BP
and 10 000 yr cal BP providing a high-resolution palaeoclimatic record for this
period in north-western Iberian margin and the adjacent continent. In
contrast, during the Holocene, mean sedimentary rate is relatively weak (10
cm kyr-1) preventing a detailed paleoclimate record for the present-day
interglacial.
X-ray analysis using SCOPIX image-processing (Migeon et al., 1999),
shows a well preserved sedimentary sequence in core MD03-2697 for the first
4. 10 m representing the last 20 000 years. Nonetheless, sediment levels of 1
cm thickness at 249 and 275 cm show the presence of Zoophycos borrows
and therefore they were not included in this work. This benthic burrowing
organism can reach a vertical extension of more than 1 m (Löwemark and
Werner, 2001; Leuschner et al., 2002) by doing a downward helicoidally
movement along a centred vertical axial shaft producing several more or less
horizontal burrows (Löwemark et al., 2004). Furthermore the successive up to
downward displacement of young material can produce important changes
in the original sediment when compared with the adjacent host sediment.
This is particularly true for the Iberian margin where foraminifera tests from
Zoophycos spreiten show younger ages of about 1 000 to 2 500 yr BP than the
adjacent sediment (Löwemark and Werner, 2001).
4. 3. 1 Stratigraphy and age model
The age model of MD03-2697 has been established by using 10 AMS
14C
dates from this core together with 1 AMS
14C
date from the twin core
MD99-2331 (42° 09’ 00 N, 09° 40’ 90 W; 2110 m depth) (Fig. IV.1 and Tab. IV.1).
These 11 dated levels were obtained at “Laboratoire de Mesure du Carbone
14” (LMC, in Saclay; France) and at Beta Analytic Inc (Beta; USA) in
199
F. Naughton, 2007
monospecific
samples
with
maxima
of
Globigerina
bulloides
or
Neogloboquadrina pachyderma (s.) abundances.
AMS
14C
dates were calibrated using CALIB Rev 5.0 program and the
"global" marine calibration dataset (marine 04.14c) (Stuiver and Reimer, 1993;
Hughen et al., 2004; Stuiver et al., 2005). We used the 95.4% (2 sigma)
confidence intervals and their relative areas under the probability curve as
well as the median probability of the probability distribution (Telford et al.,
2004) as suggested by Stuiver et al. (2005).
Some levels, such as 249 and 275 cm, were automatically rejected for
the age model reconstruction because they reflect contaminating levels of
young material introduced when the Zoophycos burrow was active
(Leuschner et al., 2002). Level 200 cm has also been considered too young
when compared with the age limits proposed by Elliot et al. (2002) for the
Heinrich 1 (H1) event. Therefore we have introduced one AMS
14C
date from
the twin core MD99-2331 for delimitate the end of H1 event in the MD03-2697.
We have also rejected level 150 cm because it has been considered too old
when compared with pollen results which place the beginning of Younger
Dryas event at 160 cm of depth.
Lab code
Beta-2131134
Beta-2131135
LMC14-003257
Beta-2131136
Beta-2131137
LMC14-003258
LMC14-003259
Beta-213138
Beta-213139
LMC14-003260
LMC14-001231
Core- depth
(cm)
MD03 269720
MD03 269740
MD03 269770
MD03 269780
MD03 2697110
MD03 2697150
MD03 2697200
MD03 2697249
MD03 2697275
MD03 2697340
MD99 2331200
Conv.
AMS 14C
age BP
Conv.
AMS 14C
age BP
(-400 yr)
error
95.4 % (2σ)
Cal BP age ranges
Cal BP
age
median
probability
G. bulloides
2880
2480
40
2501 BP:2739 BP
2656
G. bulloides
4760
4360
40
4866 BP:5198 BP
5008
G. bulloides
7435
7035
50
7783 BP:7998 BP
7895
G. bulloides
7470
7070
40
7835 BP:8014 BP
7930a
G. bulloides
9940
9540
40
10705 BP:11084 BP
10896
G. bulloides
11920
11520
60
13233 BP:13486 BP
13353 b
G. bulloides
12520
12120
70
13805 BP:14127 BP
13965 b
N. pachy (s.)
13240
12840
50
14934 BP:15424 BP
15152 a
N. pachy (s.)
13980
13580
60
15779 BP:16569 BP
16151 a
G. bulloides
15410
15010
70
18044 BP:18596 BP
18313
N. pachy (s.)
13640
13240
80
15303 BP:16099 BP
15679
Material
Tab. IV.1| Radiocarbon ages of MD03-2697 deep-sea core and one level from the twin core (MD99-2331). a
Radiocarbon dates too young or too old and b Not acceptable dating (bioturbated layers).
200
F. Naughton, 2007
54 samples covering the first 4.10 m of the MD03-2697 core were used
for pollen analysis, with a sample spacing of 1 to 10 cm. Pollen concentration
of two 1cm-thick samples (at 40 and 60 cm) was weak and therefore they
have been considered as sterile. Sample preparation technique follows de
Vernal et al. (1996) modified at the UMR CNRS 5805 EPOC (Desprat, 2005).
After chemical treatments (cold 10%, 25% and 50% HCl as well as cold
40% and 70% HF) the samples were sieved through a 10 µm nylon mesh
screen (Heusser and Stock, 1984) and mounted in bidistillate glycerine. Pollen
and spores were counted using a Zeiss Axioscope light microscope at x550
and x1250 (oil immersion) magnifications. A minimum of 100 pollen grains
excluding the over-represented Pinus grains in the Iberian margin (Naughton
et al., 2006), have been counted. At least 20 taxa and 100 grains of the
added exotic Lycopodium fern were also counted. Pollen percentages were
calculated based on the main pollen sum which excludes Pinus, aquatic
plants, spores, indeterminate and unknown pollen grains.
Pollen analysis from 0 to 2 m of the MD03-2697 deep sea core have
been previously published in Naughton et al. (2006). This previous work also
shows that pollen grains included in marine cores from western Iberia margin
reflect an integrated image of the regional vegetation of the adjacent
continent, being therefore an important tool for establishing a direct sea-land
correlation in this region.
4. 3. 3 Marine proxy analyses
4. 3. 3. 1 Ice rafted detritus (IRD) and planktonic foraminiferal
assemblages
60 and 43 levels from MD03-2697 deep sea core, with 1 to 10 cm of
sample spacing have been used for IRD and planktonic foraminifera
semiquantitative analysis, respectively. IRD and polar foraminifera analyses
from 0 to 2 m of core depth have been previously published in Naughton et
al. (2006).
Samples were washed with distilled water and wet sieved trough a 150
μm mesh screen. Planktonic foraminifera were grouped into three main
201
F. Naughton, 2007
bioclimatic assemblages: polar (Neogloboquadrina pachyderma sinistral),
subpolar (Neogloboquadrina pachyderma dextra, Globigerina bulloides and
Turborotalia quinqueloba) and warm (which includes temperate/cold,
subtropical, warm subtropical and tropical, i.e. Globorotalia scitula, G.inflata,
G.hirsuta,
Globigerina
G.truncatulinoides,
falconensis,
G.crassaformis,
G.calida,G.rubescens,
Globigerinita
G.digitata,
glutinata,
Hastigerina
aequilateralis, Orbulina universa, Globigerinoides ruber) (e.g. Bé, 1977; Ottens,
1991; Duprat, 1983).
Winter (February) and summer (August) sea-surface temperatures (SST)
were estimated by using the modern analogue technique transfer function
from the database of Pflaumann et al. (1996) improved by E. Cortijo
(Laboratoire des Sciences du Climat et de l’Environnement-LSCE, Gif-surYvette, France) and J. Duprat (UMR CNRS 5805 EPOC) on planktonic
foraminifera assemblages.
Several isotopic analyses were carried out on the first 4.10 m of MD032697 deep sea core. A total of 59 oxygen isotopic measurements were
carried out on Globigerina bulloides planktonic foraminifera and 35 on
Cibicides wuellerstorfi benthic foraminifera with a sample spacing of 1 to 10
cm and 5 to 10 cm, respectively, at the Laboratoire des Sciences du Climat et
de l’Environnement (LSCE), Gif-sur-Yvette, France. Planktonic oxygen isotopic
data from the first 2 m of MD03-2697 record has been previously published in
Naughton et al. (2006).
Each specimen has been picked up within the 250–315 µm fraction
and cleaned with distilled water. The preparation of each aliquot (4–10
specimens with 80 μg of weight) has been carried out in the Micromass
Multiprep autosampler by using an individual acid attack. The CO2 gas
extracted has been analysed against NBS 19 standard, taken as an
international reference standard. Planktonic and benthic isotopic analysis
were performed using a delta plus Finnigan at the LSCE. The mean external
reproducibility of powdered carbonate standards is ±0.05‰ for oxygen.
Results from oxygen isotopic analysis are presented versus PDB.
202
F. Naughton, 2007
4. 4 Vegetation and climate changes in north-western Iberia and
adjacent margin during the last deglaciation
High resolution continental (pollen) (Fig. IV.2) and marine proxies
(including Ice rafted detritus-IRD, planktonic foraminiferal assemblages, sea
surface temperature-SST and planktonic oxygen isotopic) were analysed
together with benthic oxygen isotopes (ice-volume indicator) from MD03-2697
deep-sea core. This has allowed the detection of sub-orbital climate
variability during the last deglaciation in north-western Iberia (Fig. IV.3):
Fig. IV.2 | Pollen diagram of MD03-2697 deep-sea core against depth. From left to right: calibrated ages
and percentages of selected pollen taxa. The stratigraphy is based on previous work by (Naughton et al.,
2006) where the Oldest Dryas represents the continental counterpart of Heinrich 1 event in the ocean.
203
F. Naughton, 2007
Fig. IV.3 | Multi-proxy record of MD03-2697 against calibrated ages. From bottom to top: percentages of
selected pollen taxa (trees: Betula, Corylus, deciduous Quercus, Pinus; Calluna and semi-desert plants:
Artemisia, Chenopodiaceae and Ephedra); Ice-rafted detritus (IRD); planktonic foraminifera associations;
Sea Surface Temperature (SST) estimates; δ18O of planktonic and benthic foraminifera and Greenland
temperatures (Sánchez Goñi et al., in prep.).
204
F. Naughton, 2007
4. 4. 1 The end of the Last Glacial maximum (LGM)
North-western Iberia was dominated by herbaceous plants with Pinus
at the end of the late pleniglacial period (Fig. IV.2). Previous works on the twin
MD99-2331 deep sea core, clearly show that this period represents the end of
the last glacial maximum (LGM) in the ocean (Naughton et al., 2006). The
expansion of heath communities (Ericaceae more Calluna) recorded in the
MD03-2697 deep sea core (Fig. IV.2) indicates moist conditions at that time as
it has been previously detected by the twin core MD99-2331 (Naughton et al.,
2006) and further south in the SU81-18 deep sea record (Turon et al., 2003).
This increase of moisture conditions in western Iberia has probably been
favoured by a more vigorous Meridional Overturning Circulation (MOC)
(Naughton et al., in prep.) which is estimated to be reduced by less than 3040% during the LGM period in comparison with the strongest reduction during
the Heinrich event 1 (H1) (McManus et al., 2004).
Sea surface conditions were warm, 11° C in winter and 16° during
summer time, corroborating previous SST reconstructions from mid-latitudes
North Atlantic deep-sea cores (Lebreiro et al., 1997; Cayre et al., 1999; Bard et
al., 2000; Chapman et al., 2000, Pailler and Bard, 2002; de Abreu et al., 2003;
de Vernal et al., 2005; Morey et al., 2005; Naughton et al., in prep.) (Fig. IV.3).
Although sea surface conditions were warmer than the previous and
subsequent Heinrich events (H2 and H1) and MOC was more vigorous,
temperate trees did not expand during the LGM interval (Naughton et al., in
prep.), though their continuous presence (~5%) suggests that scattered
woodlands survived in north-western Iberia at this time (Naughton et al.,
2006). Increasing albedo, high seasonality and low CO2 concentrations are
pointed as the major forcing mechanisms for preventing temperate tree
expansion in Iberia during this period (Naughton et al., in prep.).
4. 4. 2 The Heinrich 1 (H1)
MD03-2697 deep sea core shows a complex pattern within H1 event
(Fig. IV.3). H1 is marked by a drop in sea surface temperatures (expansion of
polar planktonic foraminifera, a decrease of summer and winter SST and
increase of planktonic δ18O values) that started at around 18 300 cal yr BP
lasting until 15 700 cal yr BP (Fig. IV.3). Although IRD presence is recorded in
205
F. Naughton, 2007
MD03-2697 deep sea core at around 18 000 cal yr BP, the maximal arrival of
Icebergs in north-western Iberian margin occurred after 16 400 cal yr BP (Fig.
IV.3).
At the same time, north-western Iberia vegetation reflects two distinct
phases. The first phase is represented by a drastic decline of Pinus forest,
testifying an important atmospheric cooling in this region (Figs. IV.2 and IV.3).
Paralleling this atmospheric cooling, a second expansion of Ericaceae and
Calluna, suggests an increase of moisture (Fig. IV.2), showing that northwestern Iberia has been influenced by different climatic conditions than those
characterizing the most continental central and eastern high-altitude sites of
Iberia, which were affected by dryer climate (see Naughton et al., 2006). In
the second phase, Pinus forest expands and semi-desert plants increase
gradually until getting a maximum synchronously with the IRD highest values
(Figs. IV.2 and IV.3).
These results are in agreement with those obtained for the late MIS 3
and during MIS 2 on the twin core MD99-2331, which suggests that each of
the typical Heinrich events (H4, H2 and H1) are marked by a first oceanic and
atmospheric extreme cold and wet episode followed by a dry and cool one
on land and in the ocean (Naughton et al., 2006; in prep.).
Recently, a study about European rivers reactivation during the last
deglaciation in the north-western French margin also shows a humid phase
preceding IRD maximal arrival (Ménot et al., 2006) (Fig. IV.1). However, this
phase has been considered as an episode preceding H1 event and, it has
been erroneously correlated with a warm continental phase. Furthermore,
previous studies on the same core, MD95-2002, have clearly showed that cold
sea surface conditions preceded IRD maximal arrival into this region (Zaragosi
et al., 2001; Auffret et al., 2002).
This strong cooling has been probably triggered by the MOC shutdown
followed by rapid oceanic and atmospheric reorganizations favouring the
transfer of atmospheric cold conditions to this region (Naughton et al., in
prep.). Because MOC shutdown prevents the moisture transfer from the North
Atlantic to Iberia, a prevailing negative mode of the NAO-like has been
proposed as the main responsible for the increase of wet conditions in northwestern Iberia during the first phase of H1 (Naughton et al., in prep.). At
206
F. Naughton, 2007
present, these prevailing atmospheric conditions could also trigger an
increase of moisture in north-western of France and generate SST increase in
north-western Atlantic (at latitudes northern than 45°-50°N) (Wanner et al.,
2001). This likely favours icebergs melting in the IRD belt and therefore, the
substantial deposition of IRD and the consequent cooling of the surface
waters in that region during the first phase of H1. So it prevents Laurentide
iceberg displacement to regions located far away from that belt, including
Iberian margin and Bay of Biscay. Prevailing negative NAO-like index also
favours the decrease of SST along north-western Iberian margin until the
Greater North Sea, passing by north-western French margin, and the slight
warmth over Greenland as shown by temperature estimates (Fig. IV.3).
On the contrary, during the second phase of H1 north-western Iberia
has been probably influenced by prevailing positive NAO-like index
(Naughton
et
al.,
in
prep.).
The
strengthening
and
the
northward
displacement of the westerlies, caused by the increase of the pressure
gradient between Açores and Iceland, favours nowadays the increase of
dryness in south-western Europe up to 47 °N and warm sea surface conditions
at mid-latitudes of the North Atlantic (20°- 40°N) (Wanner et al., 2001). This
atmospheric situation could lead the southward migration of the icebergs
until the Iberian margin during the second part of H1. During this episode of
maximum arrival of icebergs into the mid-latitudes of the North Atlantic sea
surface conditions off north-western Iberia were, paradoxically, slightly
warmer than the precedent phase. This could be explained by the influence
of prevailing positive NAO-like index along north-western Iberian margin and
Greater North Sea branch which override the cooling effect of the maximal,
although weak (Downswell et al., 1995), arrival of icebergs on this margin. The
slight warming of SST is not observed in north-western French margin probably
because the high quantity of icebergs from the British Ice Sheet arriving to this
margin. The lightening of planktonic foraminifera δ18O observed in northwestern Iberian margin within the second phase of H1 (starting at around 16
500 cal yr BP) has also been detected in other mid-latitudes deep-sea cores
such as OCE326-GGC5 (33° 42’N, 57°35’W) (McManus et al., 2004). This
decrease of δ18O values is also corroborated by the slight warming at the
North Atlantic mid-latitudes during the second phase of H1 (Fig. IV.3).
207
F. Naughton, 2007
4. 4. 3 The Bölling-Alleröd (B-A)
Deciduous Quercus expansion and herbaceous communities decline
marks the warm Bölling-Alleröd (B-A) event in north-western Iberia (Fig. IV.2).
At the same time, Iberian margin warmed as testified by the decrease of
polar planktonic foraminifera and the increase of warm and sub-polar
species. The decrease of the planktonic foraminifera δ18O values of about 1‰
PDB also suggest, beyond a decrease in salinity, an increase of sea surface
temperature during this period (Fig. IV.3).
The maximum expansion of deciduous Quercus woodlans, in northwestern Iberia, occurred synchronously with the well known meltwater pulse
1A (MWP 1A) (Fairbanks, 1989; Bard et al., 1996; 1990) and with a peak in the
Greenland temperature (Sánchez Goñi et al., in prep.), at around 14 000 cal
yr BP (Fig. IV.3).
It has been proposed by McManus et al. (2004) that the abrupt
resumption of the MOC at the beginning of B-A event together with the
increase of high latitude summer insolation accelerate the Laurentide Ice
sheet melting and trigger the MWP 1A. This ice sheet melting increased the
input of freshwater in the ocean. However, this has not produced either a
reduction or a shutdown of the conveyor belt as it would be expected.
Few hypotheses have been pointed out for explaining this controversial
situation revealed by the North Atlantic deep-sea cores. Clark et al. (2002)
proposed that meltwater pulses during the B-A were primarily originated from
Antarctica with only a small contribution of the Northern Hemisphere Ice
sheet melting while McManus et al. (2004) suggest that meltwater from the
Laurentide experienced substantial mixing with sea water before reaching
the zones of deep convection or that the site of deep-water production have
migrated further north away from the influence of the Laurentide meltwater.
Our data has shown that the timing of maximum warming during the BA period at mid-latitudes of north-western Iberia together with maxima
temperatures over Greenland coincides with that of MWP 1A. This suggests
that an atmospheric warming in the Northern Hemisphere triggered the MWP
1A event rather than the Southern Hemisphere.
208
F. Naughton, 2007
Fig. IV.4 | Long-term and small-scalle pattern of vegetation changes in north-western Iberia. From bottom
to top: benthic foraminifera δ18O; percentages of temperate trees includes (Acer, Alnus, Betula, Corylus,
Cupressaceae, deciduous Quercus, Fagus, Fraxinus excelsior-type, Pinus, Quercus ilex-type, Salix, Tilia and
Ulmus); percentages of Pinus, percentages of deciduous Quercus and summer insolation at 45° N (after
Berger, 1978).
209
F. Naughton, 2007
4. 4. 4 The Younger Dryas (YD)
Deciduous Quercus forest reduction together with the re-expansion of
semi-desert and pioneer species (Betula) mark returning cold conditions
which characterises the Younger Dryas event in north-western Iberia (Figs. IV.2
and IV.3). Synchronously with the atmospheric cooling at mid-latitudes and
over Greenland, the ocean experienced a decrease of surface temperatures
as suggested by the increase of planktonic foraminifera δ18O values (Fig.
IV.3).
These cold conditions are weaker than those characterizing H1 event (Fig.
IV.3).
Both the YD and the H1 are associated with substantial meltwater
pulses from Laurentide ice-sheet (Andrews et al., 1995). However, the YD is
only related with the reduction of the MOC while H1 with its shutdown
(McManus et al., 2004). The reduction of the MOC together with the midlatitude maximum increase of summer insolation (Fig. IV.4) could explain the
decrease rather than the complete decline of deciduous Quercus forest in
the mid- and low-altitudes of north-western Iberia. Because MOC was
activated, though in a reduced way, a weak quantity of moisture was still
transferred from the tropics to the high-latitudes of the Northern Hemisphere.
Nevertheless, Iberia (this study and see Naughton et al., 2006) and other
European regions (Watts, 1980; Lotter et al., 1992; Reille et al., 2000) were
influenced by fairly dry conditions during this period.
At the same time, the increase of trade wind strength in the Cariaco
Basin (Hughen et al., 1996) and the decrease of riverine run-off from adjacent
land masses (Peterson et al., 2000) testifies the southward migration of the
Intertropical Convergence Zone (ITCZ) which is supported by climate
modelling (Dahl et al., 2005).
Because shifts on the ITCZ are intimately connected with the NAO by
reorganizations of the Hadley cells (Marshall et al., 2001) one could think that
the southward migration of the ITCZ is accompanied by the increase of
pressure
gradient
between
Açores
and
Iceland.
This
favoured
the
intensification and northward displacement of the westerlies over Europe and
therefore an increase of dryness conditions in this region.
210
F. Naughton, 2007
4. 4. 5 The Holocene
The beginning of the Holocene is marked by the expansion of
deciduous Quercus forest and herbaceous communities decline reflecting an
increase
of
temperatures
in
north-western
Iberia
(Fig.
IV.2).
This
is
contemporaneous with a substantial warming over Greenland (Fig. IV.3) and
with mid-latitude high summer insolation (Fig. IV.4).
In the ocean, SST increases as showed by the decrease of planktonic
foraminifera δ18O values and the decrease of polar planktonic foraminifera
percentages (Fig. IV.3). MOC became more active as reflected by 231Pa/230Th
data from mid-latitude OCE326-GGC5 deep-sea core (McManus et al., 2004)
(Fig. IV.1).
The maximum spread of deciduous Quercus in north-western Iberia
representing the Holocene Thermal Maximum (HTM) in this region occurred in
the early Holocene before the 8.2 Ky event (Figs. IV.3 and IV.4). This episode
of rapid increase of atmospheric temperatures documented by the
expansion of deciduous Quercus forest in the Iberia Peninsula has also been
detected in other marine and continental pollen sequences (see Naughton
et al., 2006). Nonetheless, the HTM has been identified later, between 8000
and 4500 cal yr BP, in northern Europe (Seppä and Poska, 2004; Seppä et al.,
2005). This suggests that the Holocene maximum warming on land occurred
at different times depending of places as previously suggested for the ocean
realm by Kaufman et al. (2004).
Following the HTM in north-western Iberia, temperate forest, including
Pinus, gradually decreases and parallels the long-term trend of northern midlatitudes summer insolation (Figs. IV.3 and IV.4), suggesting that vegetation
has responded to the orbitally-induced long-term cooling that characterises
this interglacial period. This pattern has also been observed further north in
north-western France after the 8.2 Ky event (Naughton et al. submitted). This
suggests that long-term vegetation change during the mid- and lateHolocene in both regions has been triggered by orbital forcing rather than by
human impact, agreeing with what has been previously proposed by Magri
(1995).
211
F. Naughton, 2007
4. 4. 5. 1 The 8.2 k yr event
Superimposed to the long-term cooling trend of the Holocene, a
sudden decline of temperate forest, mainly involving deciduous Quercus and
Corylus trees, together with a slight increase of planktonic foraminifera δ18O
values and a decrease of SST, marks the 8.2 Ky event in north-western Iberian
margin and adjacent landmasses (Fig. IV.3). The low resolution multi-proxy
analysis for the Holocene interval in our core MD03-2697 precludes, however,
the identification of the multi-centennial cooling event between ~8600 and
8000 cal yr BP noticed by Rohling and Pälike, 2005 based in several archives
around the world and detected in the eastern North Atlantic ocean and
borderlands by Ellison et al. 2006 and Naughton et al. (submitted),
respectively.
Corylus woodlands decline at the time of the well known 8.2 ky cooling
event has been detected in several pollen sequences from central (Tinner
and Lotter, 2001; 2006) and northern Europe (Seppä and Poska, 2004; Veski et
al., 2004; Seppä et al., 2005) and in a marine pollen record from the French
margin (Naughton et al. submitted). This event was triggered by the
culminating of successive episodes that started 400 to 500 yr before this
drastic event. Laurentide Ice sheet decay and following catastrophic
outburst episodes from the lakes Agassiz and Ojibway into the Hudson Bay
(Barber et al., 1999; Teller et al., 2002; Clarke et al., 2004) allow the
introduction of freshwater into the North Atlantic (Alley et al., 1997; Clark et
al., 2001) leading to a substantial decrease of SST (Knudsen et al., 2004;
Keigwin et al., 2005, Ellison et al., 2006) and a gradual reduction of the flow
speed of the Iceland-Scotland Overflow water (ISOW), a component of the
NADW which peaked at around 8 290 yr cal BP (Ellison et al., 2006). The
impact of this drastic last event, associated with the maximum reduction of
the flow speed of the ISOW is synchronous with a temperature drop over
Greenland and Europe (Fig. IV.3).
212
F. Naughton, 2007
4. 5 Conclusion
The last deglacial period is marked in north-western Iberian margin and
adjacent landmasses by millennial scale climate variability. The end of the
Last Glacial Maximum (LGM) is marked in north-western Iberia by relatively
cold and humid conditions while mid-latitudes of the North Atlantic were
warm. We explain this apparent contradiction as the interplay of a more
vigorous Meridional Overturning Circulation (MOC) with the increasing
albedo, high seasonality and/or atmospheric CO2 drop.
Heinrich 1 is marked by a complex pattern off Iberia and in the
adjacent landmasses and represented by two main phases. The first one is
characterised by extremely cold sea surface and atmospheric conditions, a
low quantity of IRD and continental moisture. The second phase shows a high
quantity of IRD associated with continental dryness and with a slight warming
in SST and in the atmosphere although these temperatures were relatively less
cold. The extremely cold phase in north-western Iberia was probably
triggered by the MOC shutdown followed by ocean-atmosphere rapid
reorganization. We explain the shift between wet and dry conditions as the
result of changes in prevailing negative and positive NAO-like indexes,
respectively. This last mechanism can also explain the fact that the maximal
arrival of IRD is asynchronous in the different regions of the North-Atlantic
ocean. The Bölling-Alleröd (B-A) in this region is marked by a substantial
atmospheric and oceanic warming favoured by both the increase of midlatitude summer insolation and the strengthening of the MOC. The maximum
of deciduous Quercus forest expansion reflects the highest warming
conditions of north-western Iberia at 14 000 cal yr BP. This expansion is
synchronous with the Greenland temperature peak of GIS (Greenland
Interstadial)1 and with the Meltwater Pulse 1A (MWP 1A) suggesting that this
drastic melting episode must have been initiated in the Northern Hemisphere
rather that in the Southern Hemisphere.
A returning to glacial conditions characterises the Younger Dryas (YD)
event in north-western Iberian margin and in the adjacent continent. MOC
reduction, instead of shutdown, and the increase of northern mid-latitude
213
F. Naughton, 2007
summer insolation favoured the decrease rather than the complete decline
of deciduous Quercus forest in north-western Iberia. Besides the cooling,
Iberian Peninsula has been affected by substantial dryness which was
probably the result of prevailing positive NAO-like index. Oceanic and
continental warming in the early Holocene parallels Greenland temperature
curve and define the Holocene Thermal Maximum between 11 700 and 8 200
cal yr BP. Following this, the decrease of deciduous Quercus and Corylus
woodlands together with an oceanic cooling marks the 8.2 Ky event in this
region in response to the culmination of the successive episodes of the
Laurentide Ice sheet decay which enhanced the cooling over Greenland
and Europe.
The long-term trend decline of the temperate forest in north-western
Iberia during the mid- and late-Holocene parallels the gradual decrease of
northern mid-latitude summer insolation revealing that this vegetation pattern
is orbitally-induced rather than due to the human activity.
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Capítulo 5| Long-term and millennial-scale climate
variability in north-western France during the last 8 850 years
Variações climáticas de longa e pequena escala no
noroeste de França durante os últimos 8 850 anos
Variabilité climatique orbitale et sub-orbitale dans le nordouest de la France pendant les derniers 8 850 ans
The Holocene
Submitted
F. Naughton a, b, J-F. Bourillet c, M.F. Sánchez Goñi d, J-L. Turon a, J-M. Jouanneau a
a
b
Departamento de Geologia, Universidade de Lisboa, Portugal
CIFREMER,
Département
Géosciences
Marines,
Laboratoire
Environnements
Sédimentaires, Plouzané, France
dEcole
Pratique des Hautes Etudes, Environnements et Paléoenvironnements
Océaniques (UMR CNRS 5805 EPOC), Université Bordeaux 1, Av. des Facultés, 33405
Talence, France
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F. Naughton, 2007
Resumo
As variações do coberto vegetal e a estimativa dos parâmetros
climáticos obtidos numa sondagem recolhida na plataforma continental do
noroeste de França (VK03-58Bis), permitiram detectar variações climáticas
de escala orbital e sub-orbital, para os últimos 8850 anos, nesta região. O
padrão geral de diminuição gradual da temperatura de verão, marcada
pelo
declínio
contemporâneo
progressivo
do
da
floresta
arrefecimento
temperada
progressivo
da
e
húmida,
temperatura
é
na
Gronelândia, assim como, da redução dos valores de insolação de verão
nas latitudes medias até, pelo menos, 2000 anos cal BP. Ao mesmo tempo, a
diminuição da sazonalidade segue o aumento da precessão.
Entre 8739 e 8387 anos cal BP, a floresta de Corylus expande-se em
detrimento do Quercus deciduous, como resposta a uma amplificação do
contraste sazonal, resultante da expansão de gelo marinho durante o
inverno nas altas latitudes do Atlântico Norte, contrariando o padrão geral
de forçamento orbital. Este forte contraste sazonal, resulta da expulsão
drástica de água doce dos lagos de “Agassiz” e “Ojibway” e, da gradual
redução da circulação termohalina no Atlântico Norte (MOC).
Entre 8387 e 8062 anos cal BP, o súbito declínio da floresta de Corylus,
marca o evento frio “8.2 kyr”, no noroeste de França. Este episódio, foi
provavelmente produzido pela redução severa da MOC, a qual provocou
uma diminuição suplementar da temperatura de inverno, na Europa e na
Gronelândia. No entanto, o contraste sazonal permaneceu elevado durante
este evento. O forte contraste sazonal registado entre 8739 e 8062 anos cal
BP reflecte o evento multi-secular "8.6-8.0 kyr» no Atlântico Norte.
Após os estádios finais de expulsão de água dos lagos “Agassiz” e
“Ojibway”, o clima torna-se mais estável. No entanto, estão registados uma
série de ligeiros e rápidos episódios frios os quais, estão associados a um
pequeno arrefecimento no inverno e a um ligeiro aumento da precipitação.
Finalmente, a análise das associações de dinocistos e a quantificação
de gastrópodes do tipo Turritella communis permitiu-nos de detectar
variações regionais tais com: a migração para sul da zona biogeográfica
marinha Boreal entre 8739 e 8479 anos cal BP e, a abertura do canal da
mancha entre 8479 e 8387 anos cal BP.
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F. Naughton, 2007
Résumé
Les variations du couvert végétal et des paramètres climatiques, mises
en évidence à partir de l’étude d’une carotte prélevée sur la plateforme
continentale nord-ouest française (VK03-58Bis), témoignent de la variabilité
climatique orbitale et sub-orbitale des derniers 8850 ans dans cette région. A
l’échelle orbitale et jusqu’au moins 2000 ans cal BP, la reconstruction
climatique quantitative montrent un refroidissement estival qui parallélise la
régression progressive de la forêt tempérée et humide, la baisse des
températures au Groenland et la diminution de l’insolation d’été des
moyennes latitudes. La diminution de la saisonnalité détectée par le
changement
graduelle
de
la
végétation
est
compatible
avec
l’augmentation de la précession.
L’expansion de Corylus au détriment de la forêt de Quercus
caducifolié, entre 8739 et 8387 ans cal BP, est liée au fort contrast saisonnier,
amplifié par l’expansion de la banquise d’hiver dans les hautes latitudes de
l’Atlantique Nord, contrebalançant ainsi le forçage orbital. Ce fort contraste
saisonnier serait le résultat des épisodes terminaux de purges des lacs
d’Agassiz et d’Ojibway et de la réduction graduelle de la circulation
thermohaline de l’Atlantique Nord (MOC-Meridional Overturning Circulation).
Entre 8387 et 8062 ans cal BP, la régression soudaine de la forêt de
Corylus, tout en restant forte la saisonnalité, marque l’événement froid « 8.2
ka » dans le nord-ouest de la France. Cet épisode est probablement lié à la
réduction ultime et sévère de la MOC qui a provoquée une diminution
supplémentaire des températures sur l’Europe et sur le Groenland. Le fort
contraste saisonnier enregistré dans la VK03-58Bis, entre 8739 et 8062 ans cal
BP, correspond au refroidissement pluri-séculaire "8.6-8.0 ka» de l’Atlantique
Nord.
Après les derniers épisodes de purges des lacs d’Agassiz et d’Ojibway,
le climat devient plus stable. Toutefois, nos reconstructions climatiques
montrent des refroidissements d’ordre millénaire caractérisés par une légère
baisse des températures hivernales et par l’augmentation des précipitations.
De plus, l’analyse des assemblages de dinokystes et des occurrences
du gastropode Turritella communis, indiquent des changements régionaux
importants, comme la migration vers le sud de la région biogéographique
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F. Naughton, 2007
marine Boréale entre 8739-8479 ans cal BP et l’ouverture du Chenal de la
manche entre 8479 et 8387 ans cal BP.
Abstract
Vegetation and quantitative climate reconstructions from a northwestern France shelf core (VK03-58Bis) show orbital and suborbital climate
variability for the last 8850 years in this region. A long-term cooling trend in
summer temperatures, marked by gradual temperate and humid forest
decline, parallels cooling in Greenland and the decrease of mid-latitude
summer insolation reduction until at least 2000 yr cal BP. At long-term scale,
the lowering in seasonal contrast revealed by vegetation changes follows the
increase of precession.
Corylus woodlands spread at the expense of deciduous Quercus
forest, between 8739 and 8387 cal yr BP, linked with the high seasonality
conditions which, counterbalancing the long-term astronomical forcing trend,
were amplified by the north Atlantic high-latitudes winter sea-ice expansion.
High seasonality conditions resulted from the Agassiz and Ojibway final
outburst episodes and consequent gradual reduction of the MOC (Meridional
Overturning Circulation).
Between 8387-8062 cal yr BP, a sudden Corylus woodland decline
marks the 8.2 kyr cold event in north-western France probably triggered by
the severe MOC reduction leading to the additional drop in winter
temperature over Europe and Greenland. Nonetheless, seasonality remains
high during this interval. The high seasonality conditions detected in VK0358Bis between 8739-8062 cal yr BP reflects the multi-centennial-scale climate
cooling 8.6-8.0 kyr episode of the North Atlantic.
Following the Agassiz and Ojibway final outburst episodes, climate
became more stable. However, millennial scale climate cooling episodes are
recorded and characterised by weak winter cooling and increases
precipitation.
Furthermore, dinocyst
analysis and
benthic
gastropod
Turritella
communis occurrences indicate regional changes such as the southward
migration of the Boreal biogeographical zone between 8739-8479 cal yr BP
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F. Naughton, 2007
and the subsequent opening of the English Channel at around 8479-8387 cal
yr BP.
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F. Naughton, 2007
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F. Naughton, 2007
5. 1 Introduction
For a long time the Holocene interglacial has been considered a
period of stable climate. However, several studies have shown that
superimposed on the orbitally-induced long-term cooling (e.g. Kutzbach and
Gallimore, 1988; Crucifix et al., 2002; Marchal et al., 2002; Renssen et al., 2005;
Lorenz et al., 2006) sub-orbital millennial scale climate variability has affected
this interglacial (e.g. Denton and Karlén, 1973; O’Brien et al., 1995; Bond et al.,
1997; Mayewski et al., 2004).
The most extreme short-lived cold episode noticed in the Greenland
Ice cores (O’Brien et al., 1995; Alley et al., 1997; Muscheler et al., 2004), known
as the “8.2-kyr-BP event” and lasting 100-200 years, has been detected
elsewhere in several climate proxy data from the North Atlantic marine deepsea cores (Bond et al., 1997; 2001; Bianchi and McCave; 1999) and from the
European continent (e.g. Von Grafenstein et al., 1998; Klitgaard-Kristensen et
al., 1998; Nesje and Dahl, 2001; Tinner and Lotter, 2001; 2006; Baldini et al.,
2002; Magny et al., 2003; Veski et al., 2004). The causes triggering the “8.2
event” have been strongly debated over the last decade. Some authors
suggest that this event results from changes in solar activity (Denton and
Karlén, 1973; Bond et al., 2001; Van Geel et al., 2003) while others from
freshwater pulses (Von Grafenstein et al., 1999; Barber et al., 1999; Rind et al.,
2001; Alley et al., 2003).
The fact that this event is more prominent in the North Atlantic region,
that it follows two outburst flooding episodes, and that the existing similarities
between reconstructed anomaly patterns and patterns expected following a
North Atlantic freshening seem to favour the freshwater pulse mechanism as
the major trigger for the 8.2 event (Alley and Ágústsdóttir, 2005).
Two recent publications (Rohling and Pälike, 2005 and Ellison et al.,
2006) suggest that the 8.2 kyr event occurred within a long climate cooling
anomaly of multi-centennial-scale, between 8600 and 8000 years ago. This
long-lived episode has been previously noticed by a dust supply increase in
GISP2 (Mayewski et al., 1997); a decrease of sea surface temperatures (SST) in
the North Atlantic (Risebrobakken et al., 2003; Knudsen et al., 2004; Keigwin et
al., 2005) and a decrease of annual temperature from a few northern
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F. Naughton, 2007
European pollen sequences (Seppä and Poska, 2004) which the authors
usually associate with the short-lived 8.2 kyr event.
Besides the 8.2 kyr, event most of the northern European pollen
sequences from Estonia and Sweden detect a Holocene Thermal Maximum
(HTM) (Seppä and Poska, 2004; Seppä et al., 2005) between 8 000 and 4 000
cal yr BP.
So far, no studies have shown the vegetation response either to
orbitally-induced long-term cooling or to sub-orbital millennial scale climate
variability other than the extreme 8.2 kyr event in western-central Europe
during the Holocene. The aim of this study is therefore to test whether longer
and shorter-term climatic variability, involving the 8.2 kyr event and the 8.6-8.0
kyr episode, has affected western France. Towards this aim we have
performed palynological analyses (pollen and dinocysts) and pollen-derived
quantitative climate reconstructions from a shelf core VK03-58Bis retrieved in
the “Grande Vasière” of the Bay of Biscay. This core gives an integrated
image of the past regional vegetation and, therefore, the climate of western
France. This region is particularly sensitive to hydrological changes of the
North Atlantic Drift (Rahmstorf, 2002).
The Bay of Biscay presents a 300 km wide continental shelf in the northwesternmost area and becomes narrow with a steep slope further south (30
km wide) (Fig. V.1). This shelf is composed of two small and one large openshelf mud patches: the W and S Gironde shelf mud fields and the “Grande
Vasière” (Allen and Castaing, 1977). According to McCave’s classification the
“Grande Vasière” is a mid-shelf mud belt (McCave, 1972).
The “Grande Vasière” is large (more than 225 km length and 40 km
wide), located between 80 to 110 m water depth and presents an annual
mean sedimentary rate of 0.1-0.2 cm yr-1 (Lesueur et al., 2001) (Fig. V.1). Shelf
upkeep depends essentially on: a) continental supply by nepheloid layers
(Jouanneau et al., 1999; Lesueur et al., 2001); b) wave action (Pinot, 1974)
and hydrology, and c) sea level changes (Lesueur and Klingebiel, 1976). The
“Grande Vasière” rests over two sandy units and consists of a thin (few
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F. Naughton, 2007
decimetres) Holocene feature of muddy autochthon sand (Bourillet et al.,
2002). The present day spreading of sediments to the shelf is also influenced
by resuspension and redistribution of the sediments during storm episodes and
under the effects of trawling nets (Bourillet et al., 2006).
The shelf is nourished by fine grained sediments released essentially by
the Gironde and Loire rivers and to a lesser extent by the Adour, Vilaine and
Charente (Castaing and Jouaneau, 1987). The Gironde and Loire rivers have
large catchment areas (including the Massif Central and the Pyrenees zones)
recruiting pollen grains from most of the western part of France.
Indeed, previous works on world wide coastal zones with complex
fluvial systems have shown that pollen grains after being produced and
initially dispersed by the wind are mainly transported to the sea by rivers and
streams (Muller, 1959; Bottema and Van Straaten, 1966; Peck, 1973; Heusser
and Balsam, 1977; Naughton et al., in press). Experimental studies on pollen
from the French margin have shown that river systems are mainly responsible
for pollen input into the sea and that the marine pollen signature reflects an
integrated image of the regional vegetation of the adjacent continent
(Turon, 1984). Furthermore, westerly prevailing winds probably impede direct
airborne transport of pollen seaward.
Mean annual precipitation over the catchment area (PANN) varies
from 1000 mm in the westernmost part to 600 mm in the eastern zone. High
altitudinal zones, such as the Massif Central region are characterised by more
than 2200 mm of PANN while the Pyrenees vary from 2000 mm in the western
part to 1000 mm in the eastern zone. Present-day annual temperature is 13°C
in western France (data from French public Agency: “Meteo france”).
The oceanic, mild and humid climate of this region allows the
development of a temperate deciduous and warm mixed forest mainly
composed of deciduous Quercus (Q .pedunculata, Q. pubescens and Q.
sessiflora) with some scattered evergreen Quercus (Q. ilex), cork oak (Q.
suber) as well as elm (Ulmus) and ash (Fraxinus) associations. Littoral zones are
mainly composed of cluster pine (Pinus pinaster) and gorses (Ulex). There are
also beech (Fagus) and hornbeam (Carpinus) woodlands at higher altitudes.
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F. Naughton, 2007
The 2.72 m long core, VK03-58Bis, was retrieved at 96.8 m water depth
in the “Southwest-Glénan” sector of the “La Grande Vasière” mud patch
(47°36’ N and 4°08’ W) using a vibrocorer during the “Vibarmor”
oceanographic cruise (integrated in the “Défi Golfe de Gascogne” Ifremer
programme) (Fig. V.1). The “Glénan” sector is one of the end members of the
“Grande Vasière” and is composed of 3 m of sediments with high
percentages of fine material (greater than 80%).
Sedimentological analysis including micro-granulometry, calcimetry, xray analysis core using SCOPIX image-processing mode (Migeon et al., 1999)
and benthic gastropod Turritella communis counting on the VK03-58Bis shelf
core as well as the final core description were performed by Folliot (2004).
Fig. V.1 | Location of shelf core VK03-58Bis; and deep-sea core MD99-2551 (Ellison et al., 2006).
Five accelerator mass spectrometer (AMS)
14C
dates on T. communis
were obtained in the Poznan Radiocarbon Laboratory (Poland) (Bourillet et
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F. Naughton, 2007
al., 2005) indicating that the VK03-58Bis sedimentary sequence covers the last
8 850 years (Tab. V.1 and Fig. V.2). T. communis dated levels from twin cores,
VK03-58 (47°36’ N, 4°08’ W; 97.3 m water depth) and VK03-59Bis (47°38’ N,
4°09’ W; 94.6 m water depth), were correlated with that of core VK03-58Bis for
the age model construction.
All AMS
14C
dated levels were calibrated using CALIB Rev 5.0 program
and the "global" marine calibration dataset (marine 04.14c) (Stuiver and
Reimer, 1993; Hughen et al., 2004; Stuiver et al., 2005). This dataset uses the
global marine age reservoir correction (R) of 400 years. For accommodating
local effects, we have introduced the difference Δr (of about 3 years) in
reservoir age of the Bay of Arcachon (France), the closest area to our core,
as suggested by Stuiver et al. (2005). We used the 95.4% (2 sigma) confidence
intervals and their relative areas under the probability curve as well as the
median probability of the probability distribution (Telford et al., 2004) as
suggested by Stuiver et al. (2005).
Lab
code
Coredepth
(cm)
Material
POZ10166
POZ10167
POZ10168
POZ10170
POZ10171
POZ6079
POZ10172
POZ –
6077
VK03 58Bis
106
VK03 58Bis
149
VK03 58Bis
160
VK03 58Bis
177
VK03 58Bis
226
VK03 59Bis
190
VK03 59Bis
212
VK03-58
201
T.
communis
T.
communis
T.
communis
T.
communis
T.
communis
T.
communis
T.
communis
T.
communis
Conv.
AMS 14C
age BP
Conv.
AMS 14C
age BP
(-400 yr)
error
Weighted
Mean Δr
Arcachon
France
95.4 % (2σ)
Cal BP
age ranges
Cal BP age
median
probability
3820
3420
30
3
3667 BP:3865 BP
3763
7020
6620
30
3
7427 BP:7576 BP
7507
8030
7630
30
3
8391 BP:8576 BP
8479
8170
7770
30
3
8532 BP:8808 BP
8652
8240
7840
30
3
8613 BP:8938 BP
8764
7920
7520
40
3
8298 BP:8476 BP
8377
8200
7800
40
3
8567 BP:8884 BP
8696
8090
7690
50
3
8411 BP:8692 BP
8545
Tab. V.1| Radiocarbon ages from VK03-58Bis and VK03-58 and VK03-59Bis shelf cores.
5. 3. 2 Pollen and dinocyst analyses
42 and 15 samples were collected with a sample spacing of 4 to 8 cm
along the VK03-58Bis sedimentary record for pollen and dinocyst analysis,
respectively. The treatment used for palynological analysis followed the
procedure described by de Vernal et al. (1996), slightly modified at the UMR
CNRS 5805 EPOC (Desprat, 2005).
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F. Naughton, 2007
Chemical digestion using cold HCl (at 10%, 25% and 50%) and cold HF
(at 40% and 70%) were applied to eliminate carbonates and silicates. A
Lycopodium spike of known concentration was added to each sample to
calculate pollen concentrations. The residue was sieved through 10 µm nylon
mesh screens (Heusser and Stock, 1984) and mounted in bidistillate glycerine.
Pollen and cysts were identified and counted using a Zeiss microscope
with x550 and x1250 (immersion) magnifications, the last one only applied for
pollen analysis. At least 100 pollen grains (excluding Pinus, aquatic plants and
spores) and at least 15 pollen types were counted. Pinus pollen is usually overrepresented in marine deposits and therefore is often excluded from the main
sum (Heusser and Balsam, 1977; Turon, 1984). However, it is known that the
percentages of this taxa increase seaward although total pollen content
decreases (Muller, 1959; Groot and Groot, 1966; Bottema and Van Straaten,
1966; Koreneva, 1966; van der Kaars and Deckker, 2003). Because the site
location is closed to the present-day coast line we assume that Pinus pollen
percentages are not over-represented in this core and, therefore, Pinus pollen
grains have not been excluded from the main pollen sum.
Pollen percentages of each taxa were calculated based on the main
pollen
sum
that
excludes
aquatic
plants,
pteridophyte
spores
and
indeterminable pollen. 57 to 424 cysts were counted and interpreted by
comparison with modern dinoflagellate cyst distribution (de Vernal et al.,
1998; Rochon et al., 1999).
5. 3. 3 Pollen-based quantitative climate reconstruction
Quantitative climate reconstruction of north-western of France for the
last 8850 years was obtained by applying the modern analogue technique
(MAT) (Guiot et al., 1989; Guiot; 1990) to the VK03-58Bis pollen sequence. This
method is based on a modern pollen assemblage dataset including 1328
pollen spectra from Europe, Eurasia and North Africa (Peyron et al., 1998;
Peyron et al., 2005), and it selects the 5 modern pollen assemblages closest to
the fossil pollen spectra. These 5 analogues present the smallest chord
distance (Guiot, 1990) representing the best modern analogues for a given
fossil pollen spectrum and, therefore, the best samples for climate parameters
estimate. Climate parameter estimates are obtained by taking a weighted
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F. Naughton, 2007
average of the values for all selected best modern analogues which
represents the inverse of the chord distance.
Each modern analogue sample is associated with several climate
parameters which have been previously interpolated from meteorological
stations by using an Artificial Neural Network (ANN) technique (Peyron et al.,
1998). The parameters selected for climate reconstruction of north-western
France are: TANN (mean annual temperatures); PANN (mean annual
precipitation) and the difference between the temperature of the warmest
(MTWA) and the coldest (MTCO) months (seasonality). These climate
parameters are understood to play a prominent role on the distribution of the
vegetation and related pollen assemblages (Peyron et al., 2005).
5. 4 Results
5. 4. 1 Lithostratigraphy and age model
VK03-58Bis is characterised by a homogenous silt sequence marked
between 210 and 150 cm by a level containing T. communis (Fig. V.2).
Between 210 and 160 cm this T. communis community presents all the
characteristics of a biocenose: the shells are deposited in life position; both
young and adult specimens are present within the same level; they do not
present any evidence of shelf destruction by transport.
Between 160 and 150 cm, there is an increase in T. communis
abundance, and in contrast with the underlying level they are not in life
position. This indicates a drastic change in the environmental conditions
which probably resulted in their mortality.
The age obtained from the bottom of this layer in VK03-58Bis shelf core
is 7630 yr BP (8479 cal yr BP). This single drastic episode has also been
observed in the twin cores: VK03-58 dated at the bottom (7690 yr BP; 8545 cal
yr BP) and VK03-59Bis (at 4 km of distance) between 7520 and 7700 yr BP
(8377-8550 cal yr BP) (Tab. V.1). Considering the shortness of this drastic T.
communis mortality episode we can assume that this event has been
synchronous in the three cores and, therefore level 150 cm in VK03-58Bis can
be correlated with the top of that layer dated at 7520 (8377 cal yr BP) in core
VK03-59Bis.
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F. Naughton, 2007
Because there is no sedimentological evidence (no erosional surfaces
from the RX data and continuous grain size decrease) for a hiatus phase after
this drastic episode in our core, we decided to reject the date obtained for
level 149 cm which seems too young (6620 yr BP, 7507 cal yr BP) when
compared with the age limits of the Turritella layer of the twin cores.
5. 4. 2 Evolution of dinocyst assemblages
Dinocyst analysis performed in VK03-58Bis between 262 and 98 cm
shows a unique assemblage essentially composed of: Lingulodinium
machaerophorum, Operculodinium centrocarpum, and several species of
Spiniferites (Spiniferites lazus, Spiniferites bentorii, Spiniferites spp., Spiniferites
ramosus,
Spiniferites
mirabilis,
Spiniferites
membranaceus,
Spiniferites
delicatus, Spiniferites bulloideus, Spiniferites belerius, Spiniferites elongates).
Spiniferites dominates the dinocyst associations between 262 and 180
cm and is replaced by Lingulodinium machaerophorum between 180 and
160 cm.
A drastic decrease in Lingulodinium machaerophorum is detected
between 160 and 150 cm contemporaneous with the T. communis mortality
episode.
Finally, and above 150 cm, all species are replaced by Lingulodinium
machaerophorum
which
again
completely
dominates
the
dinocyst
assemblages (Fig. V.2).
5. 4. 3 Vegetation succession and quantitative climate reconstruction
Pollen analysis of the VK03-58Bis shelf core records eight main pollen
zones (numbered from the bottom to the top and prefixed by the
abbreviated sequence name VK03-58Bis) (Fig. V.2). The establishment of
these 8 pollen zones has been performed by using qualitative fluctuations of
a minimum of 2 curves of ecologically important taxa (Pons and Reille, 1986).
To delimitate chronologically each pollen zone, we have used interpolated
ages assuming a constant sedimentary rate between two consecutive dated
samples. Fig. V.3 shows the percentage curves of selected pollen taxa
plotted together with the curves of climatic parameter estimates (PANN,
MTCO, MTWA, Seasonality and TANN).
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F. Naughton, 2007
The first pollen zone (VK03-58Bis-1), 266-245 cm, (7897-7867 yr BP; 88558807 cal yr BP - extrapolated age assuming the same sedimentary rate than
that obtained between 226 and 177 cm) reflects a Pinus and deciduous
Quercus forest with Corylus and Ulmus (Fig. V.2). Quantitative climate
reconstruction shows that TANN and PANN values are 3 to 2°C and 200 mm
lower (10-11°C, 600 mm), respectively, than present day values (13°C, 800
mm) (Fig. V.3). The expansion of deciduous Quercus forest associated with
the slight spread of Corylus, Betula and Ulmus and the gradual contraction of
pine are indicated by VK03-58Bis-2 pollen zone (245-215 cm, 7867-7824 yr BP;
8807-8739 cal yr BP) (Fig. V.2). This pollen zone also suggests the presence of
scattered pockets of Acer, Fraxinus excelsior-type, Alnus and Tilia, Climatic
reconstruction estimates an increase of precipitation (150-200 mm), a
decrease of seasonality ( ΔS (summer-winter) =5°C) and a slight cooling in
summer by 4°C (Fig. V.3). In several French continental sequences such as
those from the Pyrenees and the Massif Central, the first occurrence of Tilia
has been recorded later, at the beginning of the Atlantic period (7500-5000 yr
BP; 8321-5734 cal yr BP) (Reille, 1990b; de Beaulieu et al., 1984). However,
other sequences such as that of the Soucarat in the Eastern Pyrenees records
the appearance of Tilia earlier, at around 7740±180 yr BP (8575 cal yr BP)
(Reille and Andrieu, 1994). Tilia has been also detected earlier (before 7800 yr
BP; 8700 cal yr BP) in several central-European pollen sequences such as
Soppensee and Bibersee (Switzerland), Schleinsee (Germany) (Tinner and
Lotter, 2001; 2006) and in northern European sequences such as those of
Raigastvere, Viitna, Rõuge, Ruila (Estonia) (Seppä and Poska, 2004; Veski et
al., 2004).
The next pollen zone, VK03-58Bis-3 (215-151 cm, 7824-7531 yr BP; 87398387 cal yr BP) reflects the maximum expansion of Corylus associated with the
contraction of the deciduous Quercus forest. This suggests an important
increase of seasonality between 8700 and 8200 cal yr BP that is supported by
climate estimates (Fig. V.2 and Fig. V.3). It is widely known that Corylus
competed against deciduous Quercus trees mostly during the early Holocene
in southern Europe (Tallantire, 2002). Corylus is a light-demanding tree and its
expansion is favoured by forest openings (Bradshaw and Hannon, 2004).
Furthermore, Corylus is considered to be one tolerant climate species
235
F. Naughton, 2007
supporting high seasonality conditions (Tallantire, 2002). In French continental
sequences, the maximum expansion of Corylus has been documented during
the Boreal period (9000-8000 yr BP; 10184-8866 cal yr BP) (de Beaulieu et al.,
1984; Reille and Andrieu, 1991; Reille and Lowe, 1993). Nevertheless, Buzy and
Barbazan pollen diagrams (Reille and Andrieu, 1991) as well as other
sequences from the central and eastern Pyrenees (Biscaye, Lourdes,
Freychinede, le Monge; Landos, Pinet 1, La Moulinasse 4 and Laurenti) (de
Beaulieu et al., 1984; Reille and Andrieu, 1995; Reille, 1990a; Reille and Lowe,
1993) provide evidence for a longer period of Corylus optimum extent. VK0358Bis-3 pollen zone also records a Pinus forest re-expansion. Betula and Ulmus
are consistently present and there are sporadic occurrences of Alnus, Fraxinus
excelsior-type and Tilia. Occurrences of Fagus at 8652 and 8448 cal yr BP
reflect a slight decrease in seasonality and temperatures within the period of
high seasonality that characterises VK03-58Bis-3 pollen zone.
VK03-58Bis-4 pollen zone (151-147 cm, 7531-7240 yr BP; 8387-8062 cal yr
BP) is marked by a drastic reduction of Corylus woodlands and an increase of
Pinus forest along with the maintaining of deciduous Quercus forest (Fig. V.2).
Climate estimates show that the onset of seasonality decrease coincides with
the Corylus minimum extent at around 7427 yr BP (8272 cal yr BP) (Fig. V.3).
This episode of Corylus decline has been also observed in several centralEuropean (Soppensee and Bibersee in Switzerland and Schleinsee in
Germany; Tinner and Lotter, 2001; 2006) and northern European pollen
sequences (Raigastvere, Viitna, Rõuge, Ruila in Estonia and Lake Flarken in
Sweden (Seppä and Poska, 2004; Veski et al., 2004; Seppä et al., 2005).
Corylus deflection has been interpreted as the vegetation response to the
well known 8.2 ka cooling event.
Deciduous Quercus forest attained its maximum expansion in the
following period (VK03-58Bis-5 pollen zone, 147-102 cm, 7240-3289 yr BP; 80623621 cal yr BP) suggesting a change in climate to milder (reduced
seasonality) conditions as the result of MTCO increase (Fig. V.2 and Fig. V.3).
These conditions, together with an increase of precipitation, favoured the
establishment of Alnus, Ulmus, Tilia, Fraxinus excelsior-type and Fagus trees in
western France.
236
F. Naughton, 2007
The next zone, VK03-58Bis-6 (102-55 cm, 3289-1749 yr BP; 3621–1953 cal
yr BP), indicates the slight contraction of deciduous Quercus forest, the
expansion of Fagus and the gradual increase of herbaceous plants. The slight
decrease of MTWA and the increase of MTCO lead to this vegetation
dynamic in which the Fagus spread has been probably favoured by weak
seasonality and high precipitation (Fig. V.3). In almost all the continental
French sequences, Fagus spread occurred between 4500-4000 and 2000
years BP (5000-4400 and 2150 cal yr BP) coinciding with the beginning of the
oak forest decline (Reille and Lowe, 1993; Reille and Andrieu, 1995; Reille et
al., 2000). In our VK03-58 Bis pollen record, the first occurrence of Fagus is
recorded at around 8652 and 8448 cal yr BP. Several occurrences were
detected during and after the Corylus regression episode. The beginning of a
continuous presence of Fagus occurred at around 4352 yr BP (4812 cal yr BP)
just after the Corylus regression although its maximum expansion started later
(3289 yr BP; 3621 cal yr BP). Tinner and Lotter (2001; 2006) based on pollen
analysis
from
Soppensee
and
Bibersee
(Switzerland)
and
Schleinsee
(Germany) suggest, as our sequence, that Fagus expands after the episode
of Corylus deflection, favoured by more humid summer conditions and less
extreme seasonality.
VK03-58Bis-7 pollen zone (55-24 cm, 1749-733 yr BP; 1953-852 cal yr BP)
still shows the gradual reduction of deciduous Quercus forest and the
maximum expansion of Poaceae. Fagus is still present in this pollen zone until
1061 yr BP (1207 cal yr BP) (Fig. V.2 and Fig. V.3). The continuous presence of
Cerealia type, Juglans and Castanea testifies to agricultural practices at
around 2000 years ago in western France. In the last pollen zone, VK03-58Bis-8
(upper 24 cm, last 733 yr BP; 852 cal yr BP) there is a strong increase of Pinus,
heathlands and herbaceous plants, mainly Taraxacum and Cyperaceae.
Deciduous Quercus forest decrease and Fagus virtually disappears from this
region.
237
F. Naughton, 2007
Fig. V.2 | Lithology and synthetic pollen diagram against depth (cm). From left to right: radiocarbon and
calibrated ages; lithology (after Folliot, 2004) including T. communis level (represented by small shells);
dinocyst percentages (Operculodnium centrocarpum; Total of Spiniferites and Lingulodinium
machaerophorum); pollen diagram and pollen zones.
238
F. Naughton, 2007
Fig. V.3 | Pollen diagram and quantitative pollen-based climate estimates against depth (cm). From left to
right: calibrated ages; selected pollen taxa from the synthetic pollen diagram (other deciduous trees
include: Fraxinus excelsior-type, Tilia and Ulmus); climate parameters: PANN (mean annual precipitation);
difference between the temperature of the warmest (MTWA) and the coldest (MTCO) months (seasonality)
and TANN (mean annual temperatures). Dashed lines represent maxima (bold) and minima values and the
dark line represents mean values. Grey dashed lines represent the tendency of each curve; pollen zones.
239
F. Naughton, 2007
5. 5 Climate variability in north-western France
5. 5. 1 Long-term cooling pattern and the Holocene thermal maximum
Vegetation changes and pollen-based quantitative climate estimates
permit the detection of a small-amplitude long-term pattern of summer
temperature decrease between 8850 and 2000-1000 yr cal BP.
The long-term cooling is marked by a general trend of temperate and
humid tree decline and by the increase of herbaceous plants. This long-term
cooling is characterised by the gradual decrease in the MTWA (mean
temperature of the warmest month) values (from 20.5° to 17.5° C) (Fig. V.4
and Fig. V.2) coinciding with the general trend of mid-latitude summer
insolation reduction until at least 2000 cal yr BP (Fig. V.4). The continuous
decrease of seasonality follows the gradual increase of the precessional
signal. Nonetheless the weak values in precession between 8855 and 8000 cal
yr BP surprisingly coincides with an interval of particularly high seasonality
suggesting that other mechanisms have probably amplified this precessional
signal (see below). This suggests that long-term vegetation changes in the
north-western France seem to respond directly to the Holocene orbital
induced climatic variability on which human impact on vegetation was
superimposed since at least 2000 cal yr BP. Temperate and humid forest
decrease also mimics the general decreasing trend observed in the δ18Oisotope composition of the NorthGRIP ice-core (Johnsen et al., 2001) (Fig.
V.4). These results are in agreement with the previous suggestion that forest
recession along the Holocene might be the result of natural processes rather
than a consequence of the human impact (Magri, 1995).
Previous studies on sea surface conditions of the North Atlantic and
Mediterranean regions have shown an apparent long-term cooling trend that
was driven by northern high latitudes summer insolation decreases during the
Holocene (Marchal et al., 2002; Andersen et al., 2004; Moros et al., 2004).
Several climate models also suggest an orbitally-induced mechanism as the
main forcing factor for the long-term climatic trend over the Holocene
(Kutzbach and Gallimore, 1988; Crucifix et al., 2002; Weber and Oerlemans,
2003; Renssen et al., 2005). Other authors, Lorenz et al., 2006, have compared
global alkenone-derived sea-surface temperature (SST) data with transient
climate simulations using a coupled atmosphere-ocean general circulation
240
F. Naughton, 2007
model (AOGMC) for the last 7000 yr cal BP (less instable climate period)
suggesting that mid- to late Holocene long-term SST trends were driven by
insolation changes. The general cooling trend of the Holocene starts
generally during or after the well known Holocene thermal maximum (HTM).
However, the Holocene warming that defines the HTM occurred on different
times depending of places (Kaufman et al., 2004). Several studies on both
North-Atlantic marine and Greenland ice cores detect the HTM period at the
beginning of the Holocene (Andrews and Giraudeau, 2002; Marchal et al;
2002; Duplessy et al., 2001; Kaufman et al., 2004; Knudsen et al., 2004; de
Vernal et al., 2005) while others point to a later climatic optimum (DahlJensen et al., 1998; Bauch et al., 2001; Johnsen et al., 2001; Levac et al., 2001;
Kaplan et al., 2002; Solignac et al., 2004; Kaufman et al., 2004; Keigwin et al.,
2005).
Unfortunately VK03-58Bis shelf core does not cover the entire Holocene
record. However, MTWA values were higher between 8855 and 8000 cal yr BP
than between 8000 and 1000 cal yr BP contrasting with the MTCO (mean
temperature of the coldest month) trends which show lower values during the
late early-Holocene than during the mid- and late-Holocene (Fig. V.3 and Fig.
V.4). This strong seasonal contrast between 8855 and 8000 cal yr BP likely
favoured the development of Corylus woodlands at the expense of
deciduous Quercus forest although MTWA values were high. Mild (lower
seasonality) conditions which allowed the expansion of deciduous Quercus
forest in north-western France occurred roughly between 8000 and 4000 cal
yr BP. This period has been considered as the HTM in the westernmost part of
central Europe because MTWA reconstruction shows higher values than those
from nowadays (Davis et al., 2003). Other pollen-based climate estimates
obtained from several northern European pollen sequences such as Lake
Raigastvere, Lake Viitna, Lake Ruila in Estonia and Lake Flarken in Sweden
(Seppä and Poska, 2004; Seppä et al., 2005) shows between 8000 and 4500
cal yr BP higher TANN (mean annual temperatures) values than present-day,
reflecting the HTM in those regions. In contrast, our climate estimates does not
detect either higher MTWA or TANN than present day values but reduced
seasonality between 8000 and 4500 cal yr BP. Furthermore, the continuous
presence of Fagus within this period also suggests (Tinner and Lotter, 2001)
241
F. Naughton, 2007
besides
a
weak
seasonality,
cooler
summers
and
moist
conditions.
Interestingly, during the HTM defined in central and northern Europe, the
middle latitudes of Western Europe were not submitted to particular high
temperatures precluding the identification of the HTM in this region during the
last 8850 cal yr BP.
Fig. V.4 | Correlation between vegetation changes, quantitative climate estimates, summer insolation at
45° N and precessional signal (after Berger, 1978) and δ18O-isotope composition of the NorthGRIP ice-core
(Johnsen et al., 2001) during the Holocene. Temperate and humid trees include: Acer, Alnus, Betula,
Corylus, Cupressaceae, deciduous Quercus, Fagus, Fraxinus excelsior-type, Pinus, Quercus ilex-type, Salix,
Tilia and Ulmus while Brassicaceae, Caryophyllaceae, Asteraceae (including Aster- and Anthemis- types)
and Taraxacum-type, Cyperaceae, Ericaceae and Calluna, Plantago, Poaceae and semi-desert plants
(including Chenopodiaceae, Artemisia and Ephedra) are integrated in the herbaceous plants association.
242
F. Naughton, 2007
5. 5. 2 Sub-orbital climate variability
Superimposed to the orbital induced long-term cooling pattern, pollen
analysis and quantitative climate reconstruction from VK03-58Bis shelf core
detects sub-orbital climatic variability during the last 8855 cal yr BP (Fig. V.5).
Fig. V.5 | Correlation between selected pollen taxa, quantitative climate estimates (PANN, TANN, MTCO,
MTWA and seasonality) and δ18O-isotope composition of the NorthGRIP ice-core (Johnsen et al., 2001)
during the Holocene. The 8.2 kyr event is represented by the dark grey bar which is superimposed to 8.68.0 kyr event represented by the light grey bar. Dark arrows indicate possible millennial-scale cooling
events during the Holocene.
243
F. Naughton, 2007
5. 5. 2. 1 The multi-centennial-scale climate cooling and the 8.2 ka event
The
maximum
expansion
and
subsequent
decline
of
Corylus
woodlands (between 8739-8062 cal yr BP) associated with a high seasonality
episode in north-western France occurred synchronously with the following
event succession: the last stages of the Laurentide Ice sheet decay, the
catastrophic final drainage episodes of the “glacial lakes Agassiz-Ojibway”
(Clarke et al., 2004) into the Hudson Bay, at around 8470 cal yr BP (error range
of 8160- 8740 cal yr BP; Barber et al., 1999), and the consequent 8.2 kyr event
(Teller et al., 2002; Clarke et al., 2004). The introduction of large amounts of
freshwater into the North Atlantic (Alley et al., 1997; Clark et al., 2001) triggers
an important decrease of sea surface temperatures (SST) earlier than the
recorded isotopic signal of the 8.2 kyr event in the Greenland ice cores and
lasting several centuries, between ~8900 to 8000 cal yr BP (Ellison et al., 2006).
This multi-centennial SST cooling detected by the high resolution North
Atlantic
deep-sea
core
MD99-2251
(Fig.
V.1)
occurred
roughly
contemporaneously with the climate cooling defined by Rohling and Pälike
(2005) (~8600 and 8000 cal yr BP) (Ellison et al., 2006). SST cooling (~8600 and
8000 cal yr BP) has also been observed in other regions of the North Atlantic
such as over the Laurentian Fan (Keigwin et al., 2005) and in the north of
Iceland (Knudsen et al., 2004).
The cooling and freshening of the surface ocean, that started at
around 400-500 yr before the drastic 8.2 kyr event, is linked with the beginning
of a long and gradual pattern of reduction in the flow speed of IcelandScotland Overflow water (ISOW), a component of the NADW, which attains
the slowest flow speed at around 8290 yr cal BP and lasted 200 yr,
concomitant with the 8.2 kyr event (Ellison et al., 2006). The introduction of
large amounts of freshwater favoured the reduction of the North Atlantic
Deep Water (NADW) formation (Clark et al., 2001) and the consequent
weakening of the conveyor belt (Barber et al., 1999; Rahmstorf, 2002; Renssen
et al., 2001). This mechanism has a great impact on the spread of winter seaice in the North Atlantic region playing an important role on seasonality
increase (Denton et al., 2005).
We propose that the amplified signal of seasonality in north-western
France has been driven by the final episodes of Agassiz and Ojibway
244
F. Naughton, 2007
outbursts, through the winter sea ice expansion in the high latitudes of the
North Atlantic region triggering the beginning of the maximum spread of
Corylus woodlands (at around 8739 cal yr BP).
The maximum Corylus woodlands (8739-8387 cal yr BP) expansion,
related with colder winters, is almost synchronous with the T. communis level
(8739-8479 cal yr BP) in VK03-58Bis shelf core (Fig. V.2) but also with its decline
(8479 and 8387 cal yr BP). T. communis can locally occur in fine sandy beds of
the south-western French shelf (Glemarec, 1969) although is commonly found
nowadays in the Boreal marine biogeographical zone of the north Atlantic
region, between 50 and 68 °N, further northern north than the Lusitanian
region from where VK03-58Bis core was retrieved (Fig. V.6) (Funder et al.,
2002). It is known that during the warmest phases of the early Eemian the
boreal marine zone migrated further north through the Barents and Kara sea
to the Taymyr (Funder et al., 2002). On the contrary during the early Holocene
and, in particular, between 8739-8479 cal yr BP, when the north Atlantic seaice cover was most likely extended further south as the result of the decrease
in winter temperatures, this Boreal biogeographical zone was probably
deflected several degrees further south allowing the settlement of the T.
communis off north-western France.
Between 8479 and 8387 cal yr BP a drastic environmental change
triggered the T. communis death and the decrease of the dinocyst
Lingulodinium machaerophorum (Fig. V.2). This change can not be due to
the relatively low decrease (1-3° C) of Holocene SST (Bond et al., 1997) and
salinity because both species tolerate great amplitude changes (Funder et
al., 2002; Turon, 1984; Lewis and Hallet, 1997). One regional event such as the
opening of the English Channel (Fig. V.1) (9000-7500 cal yr BP, Lambeck, 1997;
8500-8400 cal yr BP, Jiang et al., 1997; 8600-8500 cal yr BP, Gyllencreutz and
Kissel, 2006) could be the main trigger for T. communis mortality and
Lingulodinium machaerophorum decline. Indeed the opening of the English
Channel contributed to a drastic hydrological, sedimentological as well as
biological change in the north-western France (Bourillet et al., 2005) which
probably affected the benthic and planktonic communities.
245
F. Naughton, 2007
Fig. V.6 | Present day and past marine biogeographical zones in the North-East Atlantic (adapted from
Funder et al., 2002). Bold dashed lines represent the limits of the present-day marine biogeographical
zones in the North-East Atlantic; Grey dashed lines represent: a) the northward displacement of the boreal
southern limit during the early Eemian (Funder et al., 2002) and b) the southward displacement of the
boreal southern limit during the during 8.6-8.0 kyr event (this work).
The 8.2 kyr event is marked in north-western France by the drastic
episode of Corylus forest decline (8387-8062 cal yr BP) (Fig. V.5) as already
observed in central and northern Europe (Tinner and Lotter, 2001, 2006; Seppä
and Poska, 2004; Veski et al., 2004; Seppä et al., 2005). Contemporaneously
with the Corylus forest decline, PANN and MTCO decreases (of about 100
mm, 2°C) and high seasonality remains important in north-western France
(Fig. V.5). Modelling-data comparison (Wiersma and Renssen, 2005) and
pollen-based quantitative climate estimates from Europe (Davis et al., 2003)
also show a temperature reduction by at least 1°C. The drastic MOC
(Meridional Overturning Circulation) reduction at 8.2 kyr BP, associated with
slowest flow of the ISOW (Iceland-Scotland Overflow Water) (Ellison et al.,
2006), has probably amplified the signal of the Greenland isotopic record
contributing to the European temperature decrease which favoured the
decline of the Corylus trees not only in north-western France but also in
Central and Northern Europe. Decreasing seasonality following the 8.2 kyr
246
F. Naughton, 2007
event favoured deciduous Quercus expansion at the expense of Corylus
woodlands.
Furthermore, our data suggests a complex pattern of annual
precipitation in north-western France during the multi-centennial cooling that
encompasses the 8.2 kyr event (Fig. V.3 and Fig. V.5). High annual
precipitations characterise the beginning and the end phases of this cooling
episode bracketing a drier period. Lake level changes in lake Annecy reveal
the same complex pattern around the 8.2 kyr event with two high levels
separated by a low one (Magny et al., 2003). Furthermore these high lake
levels, interpreted by the authors as two episodes of high precipitation, are
associated with relatively low MTWA values (Magny et al., 2003; Magny and
Bégeot, 2004).
5. 5. 2. 2 Other possible millennial scale cooling episodes
After the final episodes of the Agassiz and Ojibway outburst flooding
climate became less instable and therefore millennial scale climatic events
are less evident during the mid- and late- Holocene. Quantitative climate
estimates from VK03-58Bis show a series of small amplitude millennial-scale
variability after the 8.2 kyr event (Fig. V.5). Cooling events are marked in most
cases by an increase of precipitation values and seasonality as well as by a
slight decrease of MTCO values. Because we only have two radiocarbon
dates levels for the last 8000 years it is difficult to correlate our events with
other well dated and world wide cooling episodes (Mayewski et al., 2004).
Nonetheless these events can probably be linked with the SST coolings
detected in the North Atlantic (Bond et al., 1997; 2001) and with some of the
ten high lake levels identified in several mid-European lacustrine records that
occurred after the 8.2 kyr event (Magny, 2004).
5. 6 Conclusions
High resolution pollen analysis and quantitative climate reconstruction
from VK03-58Bis shelf core allow the detection of a small-amplitude long-term
cooling pattern as well as millennial-scale climate variability over the last 8850
years in north-western France:
247
F. Naughton, 2007
- Both the gradual decrease of temperate and humid trees and MTWA (mean
temperature of the warmest month) values follow the general trend of
northern mid-latitude summer insolation reduction until at least 2000 cal yr BP.
The general trend of seasonality decrease follows the gradual increase of
precession;
- The high seasonality conditions of north-western France (between 8739-8062
cal yr BP) was concomitant with the multi-centennial-scale climate cooling
encompassing the 8.2 kyr event. Orbital induced colder winters were likely
amplified by the increase of winter sea ice cover in the high latitudes of the
north Atlantic as the result of the final episodes of Agassiz and Ojibway
outbursts and consequent gradual reduction of the MOC (Meridional
Overturning Circulation). This increase of seasonality favoured the spread of
Corylus woodlands at the expense of the deciduous Quercus forest between
8739-8387 cal yr BP;
- Superimposed on the multi-centennial-scale climate event an extreme
winter cooling triggered the Corylus tree decline (8387-8062 cal yr BP) in northwestern France and has been identified and related to the short-lived 8.2 kyr
cooling event. Although seasonality remains important, winter temperature
over Europe and Greenland dropped due certainly to the final drastic MOC
reduction associated with slowest flow of the ISOW (Iceland-Scotland
Overflow Water);
- We have detected a complex pattern in annual precipitation within the
multi-centennial-scale cooling (between 8739-8062 cal yr BP): a relatively dry
period in north-western France was sandwiched by two episodes of wetness;
- our study also suggests several small amplitude millennial-scale cooling after
the 8.2 kyr event marked by an increase of PANN and seasonality as well as
by a slight decrease of MTCO (mean temperature of the coldest month)
values.
Finally, high occurrence of the marine mollusc T. communis points to the
migration of the Boreal biogeographical zone several degrees further south,
between 8739-8479 cal yr BP, during the first part of the multi-centennial
cooling episode. T. communis death and the decrease of the Lingulodinium
machaerophorum dinocyst, between 8479 and 8387 cal yr BP, was certainly
triggered by the opening of the English Channel.
248
F. Naughton, 2007
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Capítulo 6| HOLOCENE CHANGES IN THE DOURO ESTUARY
(NORTHWESTERN IBERIA)
Mudanças Holocénicas no Estuário do Douro (Noroeste
Ibérico)
Changements au cours de l’Holocène dans l’estuaire du
Douro (Nord-ouest de la Péninsule Ibérique)
Journal of Coastal Research
In press. 2006
F. Naughton a, b, c , M.F. Sánchez Goñi a, T. Drago b, M.C. Freitas c, A. Oliveira d
a
b
Centro Regional de Investigação Pesqueira do Sul , Instituto Nacional de
Investigação Agrária e Pescas (INIAP) (IPIMAR-CRIPSUL), Av. 5 de Outubro, 8700-305
Olhão, Portugal
c
d
Instituto Hidrográfico (IH), Rua das Trinas, 29, 1100 Lisboa, Portugal
257
F. Naughton, 2007
Resumo
O estudo sedimentológico e palinológico de uma sondagem de 20 m
colhida na embocadura do estuário do Douro (noroeste de Portugal) mostra
uma série de variações que ocorreram ao longo do Holocénico. A presença
de uma floresta regional composta por Pinus-Quercus-Alnus, entre 10 720 e 6
530 cal BP, mostra que, o princípio do Holocénico seria caracterizado por um
clima quente e húmido. A presença e o aumento, até ao final deste período
(10 720 e 6 530 cal BP), de associações de foraminíferos marinhos
(plataforma e talude) e de equinodermes, reflecte uma subida gradual do
nível do mar.
A atenuação da subida do nível do mar e a presença de um período
de forte hidrodinamismo do rio permitiram a formação de uma barreira
cascalhenta na embocadura sul do estuário. A variação radical entre
associações polínicas provenientes da vegetação regional (composta
essencialmente por árvores), transportadas pelo rio, e espectros polínicos
provenientes
da
vegetação
local
(composto
essencialmente
por
Ericaceae/Poaceae) ocorreu entre 6 530 e 1 500 cal BP e, é contemporânea
ao estabelecimento dessa barreira de cascalho. Esta mudança de
assinatura polínica sugere que a migração de rio para norte, testemunhada
pela existência de um paleovale cujo eixo se situa a sul do presente canal,
ocorreu após 6 530 cal BP.
Résumé
L’étude sédimentologique et palynologique d’une carotte de 20 m de
long prélevée dans l’embouchure de l’estuaire du Douro (nord-ouest de
Portugal), a permis de détecter une série de variations environnementales au
cours de l’Holocène. Les résultats montrent que le début de l’Holocène (10
720-6 530 ans cal BP) est marqué par un climat chaud et humide, caractérisé
par le développement d’une forêt régionale de Pinus-Quercus-Alnus.
L’augmentation graduelle de la concentration de foraminifères benthiques,
typiques des environnements de plateforme et de talus continental, ainsi que
la présence des échinodermes à la fin de cette période suggère
l’augmentation du niveau marin entre 11 500 et 6 000 ans cal BP.
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L’atténuation de la montée du niveau marin et l’augmentation de
l’hydrodynamisme de la rivière, suggérées par cette étude indique le
développement d’une barrière de gravier dans la partie Sud de cet estuaire
a partir de 6535 ans cal BP
La variation radicale entre des associations polliniques provenant de la
végétation régionale (composées essentiellement d’arbres) et transportées
donc par la rivière, et des spectres polliniques provenant de la végétation
locale (essentiellement Ericaceae et Poaceae) est datée entre 6 530 et 1 500
ans cal BP, ce qui est contemporain de l’établissement de la barrière de
gravier. Ces résultats suggèrent que la migration vers le nord du chenal
principal de la rivière, précédemment démontrée par l’existence d’une
paléo-valley au sud de l’actuel chenal principal, a pu avoir lieu aussi tôt que
6 530 ans cal BP.
Abstract
Holocene changes are recorded by sedimentology and palynology on
a 20 m long core retrieved in the mouth of the Douro estuary (north-western
Portugal). Results show that the early Holocene (10 720-6 530 cal yr BP) was
characterised by a warm and humid climate as testified by a well-established
Pinus-Quercus-Alnus regional forest. Shelf and slope foraminifera assemblages
as well as echinoderms gradually increased towards the end of this period
reflecting the sea-level rise, which occurred between 11 500 and 6 000 cal yr
BP.
A gravel barrier developed in the southern part of the estuary as a
result of sea-level rise attenuation and strong hydrodynamism of the river. A
radical change from regional fluvially transported pollen assemblages (mainly
composed of trees) to pollen spectra derived from local vegetation (mainly
Ericaceae/Poaceae) occurred between 6 530 and 1 500 cal yr BP,
contemporaneously to the settlement of the gravel barrier. This suggests that
the northward migration of the river main channel, already testified by the
existence of a palaeovalley with its axis located southward of the present
main channel, occurred as early as 6 530 cal yr BP.
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6. 1 Introduction
Coastal ecosystems are highly productive interfaces in permanently
changing processes controlled by global factors (glacio-eustatic, climatic
and sea-level changes) and local factors, such as sand barrier dynamics or
anthropogenic impact. Only the correct interpretation of past ecological and
geomorphological changes will allow us to predict future coastline responses
to these global and local factors.
In the middle of the Holocene, during sea-level rise attenuation, the
geomorphological evolution of some coastal areas depended more on local
factors (Devoy et al., 1996; Zong, 2004), such as sediment availability or land
management practices (Devoy et al., 1996), than on major sea-level and
climate changes. Recent studies from coastal systems in southern Portugal
show that sea-level rise attenuation during the mid-Holocene is partly
responsible for sand barrier formation in coastal depressions (Bao et al., 1999).
As
a
consequence,
a
beach-barrier-lagoon
environment
has
been
established and its evolution, after 5000-4000 BP, depends essentially on local
parameters, such as barrier permeability (Freitas et al., 2002, 2003; Freitas,
1995: Freitas and Andrade, 1997).
Our knowledge of coastal evolution in northern Portugal is based
almost exclusively on studies concerning the continental shelf. These studies
reveal sea-level changes (Magalhães, 2000), coast line evolution (Dias et al.,
2000) and sedimentary dynamics (Dias, 1987; Dias et al., 2002a; 2002b; Drago,
1995; Jouanneau et al., 2002; Magalhães, 1999; Oliveira et al, 2002). However,
and up to present only few studies have been carried out on the
geomorphological evolution of the northern coastal systems in response to
global and local Holocene changes (Granja and de Groot, 1996; Granja,
1999).
The aim of this work is to fill this gap by documenting the history of the
Douro estuary during the Holocene. This estuary is repeatedly overflowed by
the Douro river, one of the most important sediment suppliers of the
Portuguese northern coast and adjacent continental shelf (Dias et al., 2002a;
2002b; Drago, 1995; Oliveira et al., 2002; Magalhães, 2000; Thouveny et al.,
2000).
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The ancient course of this river followed a northeast-southwest
direction (Carvalho and Rosa, 1988) but, at present, the river follows an eastwest orientation in the estuary zone. Some authors have tentatively proposed
that this change of flow direction resulted from the development of a sand
barrier in the mouth of the estuary (Ferreira et al., 1989). However, the
chronology of both the barrier formation and the river migration remains
unknown. The causes triggering these changes and their impact on the
estuary geomorphology are also unknown.
This is why our sedimentary and palynological approach focuses on
the southern area of the Douro estuary, in S. Paio bay. We discuss the different
geomorphological processes occurring in this area during the Holocene,
based on texture, mineralogy and composition of sediments (carbonates and
organic matter). Pollen analysis will provide, in turn, useful information on
variations of the river course over the last 10000 years.
The Douro estuary is located in the western limit of the Douro basin, in
northern Iberia (Fig. VI.1a). This is a narrow funnelled estuary, 2250 m long by
1250 m wide, partially enclosed by a sand barrier formation in the southern
area (Cabedelo), close to Vila Nova de Gaia and the palaeo-valley zone
defined by Carvalho and Rosa (1988) (Fig. VI.1b).
The lithology of the Douro basin is composed essentially of granite,
schist, gneiss and quartzite rocks (Carrington da Costa and Teixeira, 1957);
present-day climate varies from very humid in the west to semi-arid towards
the east. In the western part of the Douro basin, the wet season (OctoberMarch) covers 73% of the total annual precipitation (Loureiro et al., 1986).
The present-day vegetation of the Douro basin clearly illustrates both
Mediterranean and Atlantic influences (Pina Manique, 1957, Costa et al.,
1998, Navarro Andrés and Valle Gutiérrez, 1987). In northeastern Portugal, the
continental climate favours evergreen oaks (Quercus ilex and Quercus
suber), deciduous oaks (Quercus pyrenaica and Quercus faginea) and
Juniperus spp.. Ericaceous and Cistaceous species dominate the understory
vegetation, a consequence of intense anthropogenic activities. The
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F. Naughton, 2007
vegetation is similar in the eastern part of the Douro basin. However, cork oak
(Q. suber) is absent here. The oceanic influence is particularly important in the
northwest of the basin, where the Quercion occidentale (Quercus robur in
association with Q. suber) predominates (Braun-Blanquet et al., 1956). The
spread of both Pinus pinaster and Eucalyptus globulus has been favoured by
people. The understory vegetation is largely dominated by Ulex, in association
with heathers. The river margins are colonized by Alnus glutinosa, Fraxinus
angustifolia, Ulmus spp., Salix spp. and Populus spp. and the estuary
surroundings are dominated essentially by Poaceae and Ericaceae.
Annual river sediment charge was nearly 1.8x106 m3 under natural
conditions prior to the construction of dams in the last 40 years (0.25 x106 m3).
The mean annual draining is 22.400x106 m3, equivalent to 710 m3/s of mean
annual discharge, with a maximum of 3000 m3/s and a minimum of 50 m3/s
(Loureiro et al., 1986).
Fig. VI.1 | a) Douro estuary localisation in the Iberian Peninsula. b) Core (1, 1B and 2) and surface sampling
sites. Dark circles represent core sites and white circles surface samples sites. Dark lines represent the
palaeoisobathic curves defined by Carvalho and Rosa (1988). Dashed line represents the ancient direction
of the river main channel flow and bold dark line the present day river main channel flow. This
palaeobathymetric map shows the palaeovalley of the Douro river.
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F. Naughton, 2007
Two sediment cores (1 and 1B) were retrieved by rotary perforation in
the southern edge of Douro estuary, near the river mouth (S.Paio Bay in the
Douro estuary) (Fig. VI.1b). Another core, core 2, was retrieved in the sand
barrier, at ~500 m distance from cores 1 and 1B. Core 2 attains -39.70 m OD
(Ordnance
Datum)
Sedimentological,
and
covers
geochemical
the
and
Lateglacial/Holocene
micropalaeontological
transition.
(excluding
pollen) analyses are discussed in Drago et al. (in press).
Core 1 reached a depth of 7.90 m (-4.4 m OD), hitting a gravel unit
which stopped drilling. Coring was then moved less than 50 cm from the
original site and a new perforation (Core 1B) reach 19.70 m (-16.20 m OD) of
depth. Correlation of both cores is presented in (Fig. VI.2). The basal rock
appears between 19.70 and 17.40 m (-16.20 m/ -13.90 m OD) and is
constituted by altered granite.
Fig. VI.2 | Lithlogy of the Douro sequence. Depth is represented in Ordnance Datum which is a coordinate
system for heights above mean sea level (optometric heights). Five calibrated ages are also represented
along the two cores.
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F. Naughton, 2007
The sequence was sub-sampled for different analyses, including
sedimentary characterization (texture, mineralogy of coarser sediments,
organic matter and carbonate content), palynology and AMS radiocarbon
dating.
Five organic levels (1 cm-thick) from both cores (1 and 1B) (Fig. VI.2)
were chosen according to drastic granulometry changes to be dated by
AMS
14C
(Beta Analytic, USA). Calibration of radiocarbon dates was
effectuated by Beta Analytic by using INTCAL98 (STUIVER et al., 1998) (Tab.
VI.1).
Laboratory
Depth (m)
δ13C/δ12C
Depth (m)
Date
Calibrated age
(14C years BP)
(cal years BP) (2δ)
code
OD
Beta-164370
-0.60/-0.61
4.14-4.15
-24.4 o/oo
1110 ±40
990
Beta-154312
-1.67/-1.68
5.17-5.19
-26.2 o/oo
1580 ± 40
1500
Beta-154313
-7.14/-7.15
10.64-10.65
-26.6 o/oo
5750 ± 40
6530
Beta-174809
-8.33/-8.34
11.83-11.84
-25.2 o/oo
6050+-60
6890
Beta-154314
-13.90 /13.91
17.40-17.41
-25.6 o/oo
9490 ± 60
10720
Tab. VI.1 | Radiocarbon and calibrated dates from the site under study.
6. 3. 2 Sedimentological analyses
Sixty four samples were selected according to lithological changes
observed during core description.
Sediment texture was determined by using the Flemming classification
(Flemming,
2000).
stereomicroscope
Sand
mineralogy
observation
(x500
was
determined
magnification),
by
means
following
of
the
methodology described by Dias (1987) and Magalhães (1993, 1999).
The sand samples were separated, by dry-sieving, in five different
fractions (63-125 µm, 125-250 µm, 250-500 µm, 500-1000 µm, 1000-2000 µm). A
minimum of 300 grains was counted for each fraction. Quartz, micas,
aggregates and other terrigenous represent the mineralogical association of
265
F. Naughton, 2007
fluvial origin while bioclasts (foraminifera and echinoderms) indicate marine
origin.
Gravel study followed the methodology by Dobbkins and Folk (1970).
221 pebbles were randomly selected along the gravel unit and for each one
we measured width (l), length (L), thickness (E) and sphericity (R1) and we
determinate roundness (RK) and effective sphericity (Ee).
Organic matter was calculated by loss on ignition (Craft et al., 1991)
and carbonate content by using a Bernard calcimeter (Hulseman, 1966).
6. 3. 3 Micropalaeontological analyses
The treatment used for palynological analysis follows de Vernal et al.
(1996), slightly modified by the Département de Géologie et Oceanographie
(DGO), University Bordeaux I.
Palynological treatment consists on pollen concentration by chemical
digestion using HCl (at 10%, 25% and 50%) and HF (at 40% and 70%) to
eliminate carbonates and silicates, respectively. Samples with very high
content of organic matter were treated with KOH (at 10%). All samples were
sieved through 10 µm nylon mesh screens and the residue was mounted in
bidistillate glycerine. A Zeiss microscope with x550 and x1250 (immersion)
magnifications was used for pollen observation and counting.
Pollen identifications were achieved via comparison with specialist
atlases (Reille, 1992) as well as with the DGO pollen reference collection. At
least 350 pollen grains (excluding aquatic plants and spores) and 100
Lycopodium grains (exotic pollen introduced during the sample preparation)
were counted in each of the seventy seven samples analysed to obtain
statistically
reliable
pollen
spectra
(Rull,
1987;
Maher,
1981).
Pollen
percentages were calculated based on the main pollen sum which excludes
aquatic plants, spores, indeterminate and unknown pollen. We also
determined arboreal pollen (AP) and Non Arboreal pollen (NAP) percentages
excluding Pinus pollen from the counts. Pinus is often overrepresented in
coastal records (Heusser, 1978).
Present-day pollen percentages supply from Douro river into the study
area was also determinate to help us recognize the fossil assemblage mainly
derived from fluvial transport. This study comprised five surface (Laquasup)
266
F. Naughton, 2007
samples from the left side (southern part) of the estuary, deposited after the
Douro overflow in November 2000 (Fig. VI.1b).
6. 4. 1 Chronology
AMS14C dates obtained from the Douro sequence indicate that this
sequence covers most of the Holocene, from c. 10720 cal yr BP up to present
(Tab. VI.1).
6. 4. 2 Holocene sedimentary processes in the Douro estuary
Results from textural analyses, organic matter and carbonate content
as well as the mineralogy of sand allow the identification of three main
sedimentary units (SED2, SED3 and SED4) (Fig. VI.3). SED1 unit has been
defined in the adjacent core 2 (Drago et al., in press) as representing the
Lateglacial period.
SED2 (-13.90 m / -6.81 m OD depth), c. 10720-6530 cal yr BP, is
characterized essentially by slightly muddy sand between -13.90 m and -9.89
m and by sandy mud between -9.89 m and -6.81 m. The terrigenous content
(quartz, mica and aggregates) is higher (70-100%) than the enclosed
bioclastic grains (0-30%), suggesting that the river sediment supply was
important in the southern part of the estuary. However, carbonate content
and the increase of bioclastic fragments (benthic and planktonic foraminifers
and echinoderms) towards the top of this unit (between -9.89 m and -6.81 m)
suggest an increase of marine influence. The succession from estuarine to
shelf and slope foraminifera assemblages over this period indicates gradual
sea-level rise (Drago et al., in press) and therefore a coast line retreat that
characterizes this time period (Dias et al., 2000; Granja and de Groot, 1996).
Also, the relatively high content of organic matter, towards the top of this unit,
is probably linked to an attenuation of the hydrodynamic behaviour of the
river, favouring sandy mud deposition. The increase of micas supports this
weaker hydrodynamism which is, however, repeatedly interrupted by small
episodes of stronger hydrodynamism, as indicated by the increase of quartz
grains within the fine muddy sand layers.
267
F. Naughton, 2007
Fig. VI.3 | Lithology of the Douro sequence. Depth is represented in Ordnance Datum which is a coordinate
system for heights above mean sea level (optometric heights). Five calibrated ages are also represented
along the two cores.
268
F. Naughton, 2007
SED3 (-6.81 m / -1.67 m OD depth), accumulated c. 6530-1500 cal yr
BP, is composed by a quartzite gravel layer with weak carbonates and
organic matter values. Gravel unit presents 98-100% of ballasts with mean
grain size between 16 mm and 32 mm. Fig. VI.4a and VI.4b show the limit
bands for both high and low energy beach and river environments, based on
roundness and sphericity values plotted against the grain size, proposed by
Dobbkins and Folk (1970). Pebbles from SED3 present roundness values
between 0 and 1 (mean value=0.63) and sphericity values from 0.4 to 1
(mean value=0.69), confirming the different influences applied to these
ballasts. This suggests that the quartzite grain morphology from this unit results
from both fluvial and marine processes. Furthermore, the lithology of these
pebbles is similar to that observed nowadays in the scattered quartzite ridges
along the Douro basin, indicating that they were most likely transported by
the river. However, some of these pebbles display a composite morphology
suggesting that subsequent marine processes influenced the existing fluvial
deposit. Around 6530 cal yr BP (5750 BP), sea-level was probably similar to
present-day levels, as previously shown by Dias et al. (2000) and Granja and
de Groot (1996).
Fig. VI.4 | a) Grain roundness and b) effective sphericity of gravel pebbles plotted against particle size. Two
black lines delimit the different roundness and sphericity averages defined by Dobbkins and Folk (1970),
representing the average limit of grain characterising river and/or low and high-energy beach
environments. Dark squares represent all the measures obtained and Circle represents the value means of
all measures.
269
F. Naughton, 2007
SED4 (-1.67 m - +3.5 m), dated to post c. 1500 cal yr BP, is subdivided
into three sub-zones: SED4.a (-1.67 m / -0.1 m), SED4.b (-0.1 m / +3.43 m),
SED4.c (+3.43 m / +3.5 m) and is composed of sandy mud and slightly sandy
mud layers alternating with sand and slightly muddy sand levels. Terrigenous
content dominates all sub-units, suggesting a lack of marine supply in the
coring site. This is also indicated by the near absence of carbonate content in
this estuarine level. In SED4.a micas and high content of organic matter
essentially compose the sandy mud and slightly sandy mud layers. Sand and
slightly muddy sand levels are, in turn, constituted by quartz and low organic
matter.
SED4b is represented by quartz sand with almost no organic and
carbonate material. The mineralogy of this sand unit is similar to that of the
sand layers included in SED4a, characterized by the dominance of
terrigenous sediments.
Finally, the upper 70 cm of SED4c are formed by an alternation of sand
and muddy sand levels with relatively high values of organic matter and
micas. In comparison with the previous sedimentary units (SED2 and SED3),
SED4 reflects high terrigenous inputs and slight marine influence.
6. 4. 3 Vegetation changes versus variations in pollen catchment area
during the Holocene
Previous studies have shown that an important part of the pollen input
in most estuaries is fluvial in origin (Brush, 1989, Reille, 1990; Suc and Drivaliari,
1991; Chmura et al. 1999; Dupont and Wyputta, 2003). This must be indeed
the case of Douro estuary, since it is located in the western European coast
which is dominated by north-westerly Atlantic winds (Turon, 1984). Our pollen
diagram should, therefore, reflect plant communities that have colonised the
Douro basin over the last 10720 cal years. However, Santos et al. (2001) have
shown that fluctuations in pollen frequencies of coastal sequences are not
exclusively related to vegetation changes but can also derive from local
geomorphological changes. As a result, the pollen diagram from the Douro
coastal sequence could detect changes in pollen catchment area rather
than vegetation changes in the Douro basin.
270
F. Naughton, 2007
The Douro diagram (Fig. VI.5) is divided in three main zones. These
have been numbered from the base towards the top and prefixed by the
name LAQUA. The pollen zones are interrupted by two important pollen hiati,
associated with gravel and sand levels (at -7.03 m / - 1.70 m and 0 m / + 3.43
m) whose aerobic depositional environment prevents pollen preservation.
The first pollen zone, LAQUA1 (-13.90 m / -7.03 m), dated c. 10720-6530
cal yr BP, is marked by high frequencies of arboreal pollen (20-50%), reflecting
the dominance of a mixed Pinus-Quercus forest with Alnus in the Douro valley.
Some scattered open communities of Poaceae-Ericaceae are also present in
the regional vegetation.
LAQUA2 (-1.70 m / 0 m), c. 1500-900 cal BP, is characterized by
minimum percentages of arboreal pollen (5-10%) and maximum values of
Ericaceae, Poaceae and Asteraceae pollen. The floristic composition of this
pollen assemblage is similar to that characterizing present-day local
vegetation.
The upper centimetres of sedimentation are represented by LAQUA3.
The pollen assemblage of this zone is similar to LAQUA2, however, Pinus
percentages are higher resulting from human impact. Historical archives
indicate that pine developed in coastal areas of northern Portugal at the
beginning of the last century (Figueiral, 1995). This botanical event (input of
Pinus) allows us to infer that this zone may represent less than 100 years of
sedimentation.
Two hypotheses may explain the differences observed between
LAQUA1 and LAQUA2: a) change in regional vegetation composition, from
forest to open formation, due to strong human impact in the Douro basin
since 6530 cal yr BP; or, b) a change in pollen supply area.
The first hypothesis considers that zones LAQUA1 and LAQUA2 preserve pollen
grains that were mainly transported by the river from the vegetation
colonising the Douro basin. The second hypothesis proposes that a change in
the pollen source area occurred between LAQUA1 and LAQUA2: pollen was
essentially transported by the river from the regional forests in LAQUA1, while
in LAQUA2 pollen represents the local vegetation. This second hypothesis
implies that a change of the river trajectory occurred after 6530 cal yr BP,
decreasing the deposition of fluvial-transported pollen in the coring area.
271
F. Naughton, 2007
Fig. VI.5| Pollen diagram. From the left to the right: lithology, calibrated ages, arboreal pollen, AP (total of
arboreal pollen), Pinus, NAP (total of non-arboreal pollen), pollen of herbaceous plants and pollen zones.
272
F. Naughton, 2007
To test these hypotheses, our data was compared with other pollen
diagrams from Douro basin, covering the same time period (Peñalba, 1994;
Goméz-Lobo Rodrigues, 1993; Garcia Antón et al., 1995; Allen et al., 1996;
Peñalba et al., 1997; Von Engelbrechte, 1998; Sánchez Goñi and Hannon,
1999). These diagrams show a gradual decrease, from 6000 BP (uncalibrated)
up to the present of pollen tree percentages (in particular deciduous
Quercus), probably due to human impact and/or climate. However, total
arboreal pollen during this period is clearly much more important (AP=30-80%)
in those diagrams than in LAQUA2 zone (AP=5-10%) (Fig VI.5). Therefore, a
vegetation change in the Douro basin from forest to open communities can
not explain the difference observed in the pollen record before and after
6530 cal yr BP.
The second hypothesis, involving a shift in the pollen source area,
seems the most likely explanation. To verify this hypothesis, surface sediments
were collected in the Douro estuary, after the November 2000 overflow, in
order to identify the type and frequencies of present-day pollen taxa,
transported by the Douro and incorporated in the sediments (Fig VI.6). This
surface sediment study shows that arboreal pollen percentages are high
(AP=20-30%), contrary to what is observed in LAQUA2.
Furthermore, the comparison between the modern (LAQUASUP) and
fossil spectra (LAQUA1-3) reflects a similarity between LAQUASUP and
LAQUA1, which means that nowadays, after an overflow, the river transports
a great amount of arboreal pollen from the entire Douro basin (AP=20-30%),
and this despite present-day anthropisation. It is possible that from c. 1500 cal
yr BP onwards the river might have had less importance on pollen supply to
the coring site, and this is probably related to a northward river main channel
migration, between 6530 – 1500 cal yr BP. A palaeo-valley with its axis located
south of the present main channel testifies to this northward deflection of the
river flow (Carvalho and Rosa, 1988) (Fig VI.1b).
273
F. Naughton, 2007
Fig. VI.6| LAQUASUP: surface samples sites and pollen percentages.
274
F. Naughton, 2007
6. 4. 4 Geomorphological changes in the Douro estuary during the
Holocene
10720 – 6530 cal yr BP (Fig. VI.7a)
The first part of this period is marked by the prevalence of the fluvial
influence as shown by the presence of terrigenous components, the almost
absence of foramifera and echinoderms and the presence of regional fluvialtransported pollen.
Concerning the geomorphology, the present day estuarine zone was
an ancient river course within a northeast-southwest incised valley (Fig. VI.7a),
as also shown by the palaeobathymetric map inferred by Carvalho and Rosa
(1988) (Fig. VI.1b).
The relatively warm and humid climate which characterizes the
beginning of the Holocene, favoured the development Pinus-Quercus forest
with Alnus, as reflected by their percentages contribution to our pollen
spectra during (10720 – 6530 cal yr BP).
The second part of this interval (10720 – 6530 cal yr BP), is characterized
by an increase in marine microfossils towards the top and the presence of
fluvial-transported pollen indicating a period of mixed fluvial and marine
influences in the area, a likely result of sea-level rise and coast line recession.
This is in agreement with the sea-level rise and coast line recession detected
for the NW of Iberia by Dias et al. (2000) and Granja and de Groot (1996).
6530 – 1500 cal yr BP (Fig. VI.7b)
During this period, the accumulation of 5 m of gravel suggests drastic
environmental changes probably due to more frequent torrential rains.
This gravel unit, essentially constituted by pebbles derived from
quartzite outcrops in the Douro basin that exhibit a mixed morphology,
showing that they had been initially transported by the river waters until the
core site and then, reworked by the sea.
At this time, the reduction in the rate of sea-level rise (Dias et al., 2000)
favours sediment deposition in the estuary. This, in turn, contributed to the
formation of a gravel barrier and its northward growth triggered the migration
of the Douro main channel to the north (Fig. VI.7b).
275
F. Naughton, 2007
This river main channel migration is suggested by the change in pollen
origin, from regional fluvially-transported pollen to local pollen vegetation.
Following the gravel barrier formation, local factors seem to play a major role
in the geomorphological development of the Douro estuary.
Sand barriers development has also been recorded further south, in
coastal ecosystems such as Albufeira, Santo André and Melides (Freitas et al.,
2002; 2003).
1500 cal yr BP – present-day (Fig. VI.7c)
The development of the gravel barrier in this area precluded the arrival
of marine material over the last 1500 years, as demonstrated by the lack of
carbonates and bioclast fragments in the estuarine sediments.
The local pollen signature recorded during this period suggests the
weak influence of the Douro river in the southern part of the estuary.
However, after an overflow, the study area receives some detrital
material and pollen grains from the hydrographical basin as showed by the
LAQUASUP assemblages.
The establishment of the gravel barrier in the southern area of the
Douro mouth and river main channel migration to the north had contributed
to the enlargement of the estuarine zone and settlement of present-day
geomorphology (Fig. VI.7b and VI.7c).
Fig. VI.7| Conceptual model of the Holocene geomorphological evolution of the Douro estuary: a)
between 10720 and 6530 cal yr BP, b) from 6530 to 1500 cal yr BP and c) the last 1500 years.
276
F. Naughton, 2007
6. 5 CONCLUSIONS
The multiproxy study of core 1 and 1B from the Douro estuary
documents the Holocene geomorphological history of this coastal area and
identifies the major factors explaining the evolution of this system.
Between 10720 and 6530 cal yr BP, changes in geomorphology of the
Douro estuary were essentially influenced by the global sea-level rise. Gradual
increase of marine influences in the estuary indicates this coastline recession.
Later on, a shift in pollen record shows a change in pollen source area due to
a northward migration of the Douro main channel between 6530 and 1500
cal yr BP. This was likely caused by the formation of a gravel barrier as the
result of both sea-level rise attenuation and high fluvial hydrodynamism
caused by the increase of torrential rains. Local characteristics become a
major agent controlling geomorphology evolution of the Douro estuary,
during the upper part of the Holocene.
This study further highlights the potential of pollen analysis as a proxy to
reconstruct geomorphological changes in coastal environments.
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282
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Capítulo 7| Conclusions and perspectives
In order to deeply characterise and understand the response of
vegetation and climate of south-western Europe to the climate variability of
the North Atlantic over the last 30 000 years we have carried out a high
resolution multi-proxy analysis (pollen, planktonic foraminifera assemblages,
planktonic and benthic foraminifera δ18O measurements and Uk37 sea surface
temperature reconstruction) of two marine deep-sea cores (MD99-2331 and
MD03-2697) retrieved in the Galician margin (north-western Iberia). These
marine deep-sea cores located at ~100 Km from the coast preserve a high
quantity of pollen grains from the adjacent landmasses together with several
marine climate indicators and an ice-volume proxy. This has allowed us to
establish a direct correlation between marine and terrestrial proxies and to
identify possible leads and lags in the response of the different Earth reservoirs
to a given climatic change.
Before starting this paleoclimatic study we have verified whether the
pollen preserved in the Iberian margin sediments represent one integrated
image of the regional vegetation colonising the adjacent landmasses. We
have compared modern pollen assemblages from the continent with those
from marine sediments. For this, we have analysed pollen content from
several sedimentary modern samples retrieved in south (Mediterranean
region) and in north-western (Atlantic region) Iberia following two transects
including the estuary, shelf and slope. The resulting marine and coastal
modern pollen signatures have been compared with several present-day
continental samples including mosses, surface sediments and top of peat
bog and lake sequences available in the European Pollen Database.
This comparison showed that the pollen signature from the Iberian
margin is similar to that of the Iberian terrestrial deposits, and, in particular, to
that of the estuarine samples which recruit pollen from the vegetation
colonising the adjacent hydrographic basins. Therefore, western Iberian
margin pollen spectra reflect an integrated image of the regional vegetation
of the adjacent continent. Our study has also showed that marine pollen
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F. Naughton, 2007
spectra clearly discriminate both the Mediterranean and the Atlantic plant
communities colonising southern and northern Iberian Peninsula, respectively.
Furthermore, it also demonstrate that Pinus pollen are as elsewhere
overrepresented in marine sediments indicating that this pollen must not be
included in the main pollen sum which is usually used for calculate pollen
percentages. Further, this work has allowed us to understand the mechanisms
involved in pollen transfer from Iberia to the ocean.
The western Iberian margin is at present influenced by north-western
prevailing winds which preclude substantial direct airborne transport of pollen
seaward. On the other hand, this margin is closed to several important
hydrographic basins (Tagus, Sado, Douro and Minho) favouring pollen
seaward transfer mainly by fluvial transport. Results from Total Pollen
Concentration (TPC) together with proposed conceptual models of fine
particle dynamics in the Iberian margin have allowed us to establish the
present-day pattern of pollen dispersion in this region. In north-western Iberian
margin, pollen grains released by Douro followed by Minho rivers are
enclosed in nepheloid layers and transported to the shelf until getting
blocked by the rocky outcrops. In winter, during downwelling conditions
pollen grains are then transported polewards, firstly deposited in the Douro
mud patch (S-N direction) then in the Galicia mud patch, and finally they
flow westward to the deep-sea. Only small quantities of pollen grains can be
transported directly to the outer shelf and upper slope under extreme stormy
events. In summer, under upwelling conditions, pollen transfer to the slope
must be restricted to offshore filaments. In the southern Iberian margin, pollen
grains released by the Tagus and to a lesser extent by the Sado river, are
partially deposited in the shelf and transported to the south and seaward by
littoral and oceanic currents probably through the southern canyons during
upwelling conditions.
Besides the verification of the reliability of Iberian margin pollen signatures, our
study reveals that marine pollen sequences from western Iberian margin are a
powerful tool to accurately reconstruct the vegetation response to oceanic
and atmospheric climate changes within a reliable chronological framework.
The comparison of the high resolution pollen analysis from the Galician margin
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F. Naughton, 2007
composite core (MD99-2331 and MD03-2697) with other marine and terrestrial
pollen sequences has allowed us to document vegetation changes in the
Iberian Peninsula over the last 25 000 years. It also shows that our Galician
margin sequence mainly represent the vegetation of low- and mid-altitude
zones. Further, the direct correlation between marine proxies such as IRD
content,
planktonic
polar
foraminifera
percentages
and
planktonic
foraminifera δ18O with pollen from north-western Iberia reveals that
vegetation of north-western Iberia has responded synchronously to the
climatic variability detected elsewhere in the North Atlantic realm during the
Heinrich events 2 and 1 (H2 and H1) (26 000-24 380 cal yr BP e 15 900-18 500
cal yr BP), the Last Glacial Maximum (LGM) (24 300 and 18 500 cal yr BP), the
Bölling Alleröd (B-A) warm period (13 200-15 900 cal yr BP) and the Younger
Dryas (YD) cold event (11 600-13 200 cal yr BP).
Vegetation response to the well known Heinrich events 2 and 1 (H2
and H1) is complex and characterised by two vegetation phases at low and
mid-altitudes of north-western Iberia. The beginning of each Heinrich event is
marked on land by an important pine forest reduction and the expansion of
heathers suggesting that this first phase was cold and wet. Pinus forest
expansion characterising the second phase of each Heinrich event indicates
a less cold episode associated with an increase of dryness as suggested by
the development of semi-desert associations. We have also demonstrated
that the H1 event is the marine equivalent of the Oldest Dryas in the
continent.
The Last Glacial Maximum (LGM), bracketed by H2 and H1 events, was
characterised by the expansion of Pinus in an herbaceous-dominant
environment along with scattered pockets of deciduous trees (Quercus,
Corylus and Alnus) in north-western Iberia. This suggested that not only
southern Iberia but also northern Iberia acted as a refugium zone for
temperate trees, though at a smaller scale. Furthermore, western Iberia was
characterised by prevailing wet conditions during the LGM as suggested by
the expansion of heath communities (Ericaceae including Calluna) in both
north-western and south-western regions.
285
F. Naughton, 2007
The Bölling-Allerød interstadial in our Galician sequence shows a more
rapid and strong expansion of deciduous Quercus at low- and mid-altitude
zones than in the high altitude sites of north-western Iberia. This indicates that
the temperate forest from low and mid-altitudes of north-western Iberia
together with that from southern Iberia region responded more rapidly to the
climate variability of the North Atlantic during the B-A interstadial than in the
northern high altitudes of Iberia.
The Younger Dryas cold event was characterised at low- and midaltitude of north-western Iberia by an increase of pioneer species (mainly
Betula), grasses and semi-desert associations (Artemisia and Ephedra) at the
expense of the temperate forest. However, the vegetation reversal
characterizing the Younger Dryas event was less marked in those regions than
in the high altitude sites of northern Iberia certainly due to the maximum
expansion of deciduous forest during the previous B-A interstadial and a
mitigated decrease in temperatures in the lower areas.
The response of deciduous forest to the climate improvement that
characterised the onset of the Holocene at low and mid altitudes of northwestern Iberia seems to lead those observed in the high altitude sites,
although the same succession of trees (Juniperus, Betula, Pinus Quercus,
Corylus and Alnus) was observed in all northern regions. Indeed, this
vegetation succession started during the Younger Dryas in our Galician core
while in the high altitudes of northern Iberia at the beginning of the Holocene.
In southern Iberia deciduous Quercus expansion occurred as in the mid- and
low-altitudes of north-western Iberia rapidly most likely as the result of the
higher density of refugia for temperate trees in these zones during the LGM.
The direct correlation between terrestrial and marine climatic
indicators enables us to tackle the mechanisms responsible for the complex
pattern left by Heinrich events in north-western Iberia which corresponds to
that previously observed in the mid-latitudes of the North Atlantic. Two main
vegetation phases in north-western Iberia were linked to the complex pattern
left by the typical Heinrich events (H4, H2 and H1) in the Iberian margin. The
first phase was marked by a drop of Sea Surface Temperature (SST) (indicated
by the planktonic foraminifera assemblages and δ18O analyses together with
Uk37 measurements) the virtual absence of icebergs in north-western Iberian
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F. Naughton, 2007
margin and the strong cooling of the adjacent continent revealed by the
Pinus forest decline. This first phase was also characterised by an increase of
moisture conditions in Iberia as showed by the Calluna expansion in concert
with the increase of total pollen concentration in MD99-2331 record. The
second phase, associated with the maximal arrival of icebergs into this
Iberian margin, was characterised by less cold sea surface and atmospheric
conditions (Pinus forest development) and by an increase of dryness
identified by the expansion of semi-desert plants. The impact of H3 in northwestern Iberia was peculiar and associated with prevailing wet conditions
over almost the entire event. Nonetheless, a substantial drop in SST and
atmospheric temperatures (Pinus forest decline) preceded the episode of
maximal icebergs discharges into the Iberian margin as during H4, H2 and H1
events.
We have proposed two main mechanisms underlying this complex
pattern. The extreme atmospheric cold conditions indicated by the
substantial decline of Pinus forest in north-western Iberia during Heinrich
events 4, 3, 2 and 1 were probably caused by the introduction of large
amounts of freshwater via Northern hemisphere icebergs drifting and
consequent melting triggering a shutdown of the Atlantic Meridional
Overturning Circulation (MOC). This MOC shutdown was followed by oceanatmosphere rapid reorganizations favouring the fast transfer of cold
conditions into north-western Iberian Peninsula. Superimposed to this
oceanographic mechanism, changes similar to those of prevailing (negative
to positive) North Atlantic Oscillation (NAO) index seems to have played a
crucial role for explaining this complex pattern. Indeed during the first phase,
prevailing negative mode of NAO-like index likely triggered the increase of
winter precipitation in Iberia and enhanced river flow favouring the seaward
pollen transfer. These prevailing conditions generated a SST increase in northwestern Atlantic (at latitudes northern than 45°-50°N) favouring iceberg
melting in the IRD belt preventing their southern migration to the midlatitudes. This icebergs melting produced then a drop in sea surface
conditions in the IRD belt region. This prevailing NAO-like negative mode also
produced the drop of SST along the branch between north-western Iberian
margin and the Greater North Sea amplifying the sea surface cooling in
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F. Naughton, 2007
north-western Iberian margin. During the second phase of the typical Heinrich
events (H4, H2 and H1) the change to prevailing positive mode of the NAO–
like index triggered the strengthening and northward displacement of the
westerlies favouring the increase of dryness in Europe including Iberia
Peninsula. These prevailing conditions produced relatively warm sea surface
conditions at mid-latitudes of the North Atlantic (~ 20°- 40°N) favouring the
southward migration of the icebergs to the mid-latitudes sites including
western Iberian margin. During this episode of maximum arrival of icebergs
into the mid-latitudes of the North Atlantic, sea surface conditions off northwestern Iberia were slightly warmer than the precedent phase showing the
influence of prevailing positive NAO-like index along north-western Iberian
margin and Greater North Sea branch. The prevailing wet conditions during
the atypical H3 in north-western Iberia probably resulted from the maintaining
of reduced westerlies in this region.
The land-sea direct correlation has revealed that during the LGM
period temperate tree expansion was largely reduced when compared with
the previous late MIS 3 D-O interstadials although sea surface conditions were
similar and relatively high. We have proposed three mechanisms for
explaining the decoupling between the ocean surface temperatures and the
temperate forest development: a) albedo increase which enhanced the
cooling produced by low summer insolation in the Northern Hemisphere, b)
high seasonality as a result of substantial winter cooling, and c) weak CO2
concentration. However, western Iberia have been influenced by relatively
wet conditions during the LGM likely as the response to the strengthening of
the Meridional Overturning Circulation (MOC) which was more vigorous than
during the bracketing Heinrich events.
Besides the last glacial period, we were also interested to know about
the millennial-scale climatic oscillations at mid-latitudes of the North Atlantic
and in north-western Iberia which occurred during the last deglaciation
(19 500-7000 yr cal BP), a period characterised by the increase of summer
insolation and the substantial ice volume reduction in the northern high
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F. Naughton, 2007
latitudes. This period encompasses the end of the LGM, the H1, Bölling-Alleröd
(B-A) interstadial, Younger Dryas and early Holocene (11 500-8 000 cal yr BP).
The substantial atmospheric (deciduous Quercus expansion) and seasurface warming following H1 was probably caused by both the increase of
mid-latitude summer insolation and the strengthening of the MOC. Maximum
expansion of deciduous Quercus at around 14 000 cal yr BP occurred
synchronously with the Greenland temperature peak of GIS (Greenland
Interstadial) 1 and with the Meltwater Pulse 1A (MWP 1A). This suggests that
the severe warming of the Northern Hemisphere could be the trigger
mechanism for the drastic melting episode (MWP 1A) of the Laurentide Ice
sheet implying that this event has been probably initiated in the Northern
rather than in the Southern Hemisphere.
Subsequent decrease of deciduous Quercus forest in north-western
Iberia was contemporaneous with sea surface cooling at the mid-latitudes of
the North Atlantic and reflects the Younger Dryas (YD) event in that region.
MOC reduction, instead of shutdown, and the increase of northern midlatitude summer insolation favoured the decrease rather than the complete
decline of deciduous Quercus forest, as it is the case during H1, in northwestern Iberia. Besides the cooling, the expansion of semi-desert plants
reflected an increase of dryness in Iberia which was likely the result of
prevailing positive mode of the NAO-like index in the North Atlantic realm.
Concomitant with high values in the Northern mid-latitudes summer
insolation we have detected the Holocene Thermal Maximum (HTM) in northwestern Iberia between 11 700 and 8 200 cal yr BP. At around 8 200 cal yr BP a
slight drop in SST of the eastern North Atlantic mid-latitudes together with a
sudden decrease of deciduous Quercus and Corylus forest in north-western
Iberia marked the 8.2 Kyr event in those regions. This event reflects the
culmination of the successive episodes associated with the Laurentide Ice
sheet decay which enhanced the cooling over Greenland and Europe.
Following the 8.2 ky event the gradual decrease of temperate forest
paralleled mid-latitudes summer insolation decrease showing that this longterm trend of forest reduction was influenced by orbitally-induced cooling
rather than by human impact.
289
F. Naughton, 2007
The low resolution multi-proxy analysis for the Holocene interval in
MD03-2697 deep-sea core has precluded, however, the identification of the
millennial scale climatic variability of the Holocene in the mid-latitudes of the
North Atlantic and in particular, the multi-centennial cooling event between
~8 600 and 8 000 cal yr BP detected elsewhere. For filling this gap, we have
analysed the pollen content of a 2.72 m long core, Vk03-58Bis, retrieved in
north-western French shelf (47°36’ N and 4°08’ W). Vegetation and
quantitative climatic reconstructions from this core have shown orbital and
suborbital climate variability over the last 8 850 years in north-western France.
Gradual temperate forest decline together with the decrease of MTWA
(mean temperature of the warmest month) values parallels until at least 2 000
yr cal BP the steady long-term cooling of Greenland and the reduction of
mid-latitude summer insolation. Also, the gradual decrease of seasonality
followed, as it was expected, the increase of precession. Counterbalancing
this long-term astronomical forcing trend, Corylus woodlands spread at the
expense of deciduous Quercus forest, between 8 739 and 8 387 cal yr BP,
reflecting particular high seasonality conditions as the result of the
amplification produced by the expansion of sea-ice in North Atlantic highlatitudes during winter. This increase of winter sea ice cover was triggered by
the final episodes of Agassiz and Ojibway outbursts and consequent gradual
reduction of the MOC. A sudden Corylus woodland decline in north-western
France, between 8 387-8 062 cal yr BP, already noticed in other European
regions, marked the 8.2 kyr cold event. This extremely cold episode was
probably triggered by the final drastic MOC reduction associated with
slowest flow speed of the ISOW (Iceland-Scotland Overflow Water) leading to
an additional drop in winter temperature over Europe and Greenland.
Nonetheless, seasonality remained high during this interval. Indeed the high
seasonality conditions detected in VK03-58Bis between 8 739-8 062 cal yr BP
reflected therefore the multi-centennial-scale climate cooling 8.6-8.0 kyr
episode of the North Atlantic.
Following the Agassiz and Ojibway final outburst episodes, climate
became more stable. However, millennial scale climate cooling episodes are
recorded and characterised by weak winter cooling and increases in
precipitation and seasonality.
290
F. Naughton, 2007
Besides pollen, dinocyst analysis and benthic gastropod Turritella
communis occurrences in VK03-58Bis indicated regional changes such as: a)
the southward migration of the Boreal biogeographical zone between 8 739-8
479 cal yr BP as the result of the southern extension of the North Atlantic seaice cover triggered by the decrease of winter temperatures which favoured
the settlement of the T. communis off north-western France; and b) the
subsequent opening of the English Channel at around 8 479-8 387 cal yr BP
which produced a drastic environmental change triggering the T. communis
death and the decrease of the dinocyst Lingulodinium machaerophorum.
Since the last deglaciation, climate variability has a great impact on
coastal systems evolution. Indeed the increases of high latitude summer
insolation leading to the decrease of ice volume favoured the global sealevel rise and consequent changes in coastal areas. In order to understand
the impact of global forcing factors such as climate and sea-level changes
on the geomorphology of north-western Iberia coastal systems, we have
applied a multi-proxy study of two cores retrieved in the mouth of the Douro
estuary (north-western Portugal).
The early to mid-Holocene (between 10 720 and 6 530 cal yr BP) was
characterised by a well-established Pinus-Quercus-Alnus forest reflecting
warm and moist conditions in the Douro basin. A gradual increase of shelf
and slope benthic foraminifera assemblages together with echinoderms
reflect a steady sea-level rise. Pollen grains from the regional vegetation were
transported by the river along the ancient river main channel and deposited
in the study area which was at that time a river margin.
At around 6 530 cal yr BP a drastic environmental change is marked by
the deposition of 5 m of gravel in the southern part of the estuary as the result
of high hydrodynamism of the river and sea-level rise attenuation. The
settlement of this gravel barrier between 6530 and 1500 cal yr BP leads to the
northward migration of the river main channel preventing fluvially transported
pollen grains from the regional vegetation of the Douro basin to be deposited
in the study area after 1500 cal yr BP. Indeed, the radical change from
regional fluvially transported pollen assemblages (mainly composed of trees)
to pollen spectra derived from local vegetation (mainly Ericaceae/Poaceae)
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F. Naughton, 2007
occurred contemporaneously with the settlement of this gravel barrier. The
establishment of the gravel barrier in the southern area of the Douro mouth
and river main channel migration to the north had contributed to the
enlargement of the estuarine zone. Following this, after 1500 cal yr BP, local
factors seemed to have played a major role in the geomorphological
development of the Douro estuary.
Perspectives
After this work, several questions are still unsolved….
It will be necessary to carry out some high resolution multi-proxy study
of the last 30 000 cal yr BP in the mid-latitudes of the western North Atlantic
and at higher latitudes than Iberian margin in the eastern and western North
Atlantic to:
- confirm the impact of the atmospheric mechanism linked to changes
in the prevailing negative to positive mode of NAO-like index, on Heinrich
events;
- understand if the extreme warming detected in north-western Iberia
and Greenland during the Melwater pulse 1A occurred in other North Atlantic
regions to investigate the Northern Hemisphere origin of this last event;
-to detect the multi-centennial episode between 8.6 and 8.0 ky BP
including the short-lived 8.2 ky event, the Holocene Thermal Maximum and
the millennial-scale climate variability of the mid- and late-Holocene in the
North Atlantic.
Finally, it would be also of extreme importance to correlate a multiproxy marine deep-sea core of north-western Iberia representing the last
glacial and interglacial transition (LGIT) with an estuarine core from northwestern of Portugal, in order to understand the impact of the North Atlantic
climate variability on the evolution of that coastal system during this
transitional period.
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F. Naughton, 2007
Conclusões
No intuito de caracterizar e compreender a resposta da vegetação e
do clima do sudoeste da Europa à variabilidade climática que caracteriza a
região do Atlântico Norte, durante os últimos 30 000 anos, foi efectuada uma
análise de proxies múltiplos com alta resolução temporal (pólen, associações
de
foraminíferos planctónicos, δ18O
de foraminíferos planctónicos e
bentónicos e reconstrução da temperatura superficial do oceano baseada
na análise de alcanonas) em duas sondagens marinhas (MD99-2331 e MD032697) recolhidas na margem Galega (noroeste da margem Ibérica). Estas
sondagens foram recolhidas a cerca de 100 Km da costa e possuem uma
grande abudância de grãos de pólen provenientes do continente
adjacente, assim como uma série de indicadores paleoclimáticos marinhos e
um indicador do volume de gelo acumulado nos pólos. A presença de
diferentes tipos de indicadores permite o estabelecimento de uma
correlação directa entre os diferentes sub-sistemas climáticos assim como
identificar possíveis sincronias e assincronias na resposta dos mesmos a um
dado evento climático.
Antes de iniciar este estudo paleoclimático, foi necessário verificar se
os grãos de pólen preservados nos sedimentos da margem ibérica
representavam de facto uma imagem integral da vegetação que coloniza o
continente adjacente. Desta forma, foi efectuada a análise polínica em
várias amostras superficiais colhidas no sudoeste (região Mediterrânica) e
noroeste (região Atlântica) da margem Ibérica ao longo de um trajecto
transversal à actual linha de costa incluindo estuários, plataforma e talude
continental. Estas amostras foram comparados com as assinaturas polínicas
continentais actuais provenientes da base de dados polínicos da Europa
(European Pollen Database).
Neste estudo comparativo foi possível verificar que as assinaturas
polínicas da margem ibérica são semelhantes àquelas detectadas nas
amostras continentais. De facto, as assinaturas polínicas da margem Ibérica
são particularmente semelhantes àquelas obtidas para as zonas estuarinas
as quais recrutam os grãos de pólen provenientes da vegetação que
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F. Naughton, 2007
coloniza as bacias hidrográficas adjacentes. Os resultados obtidos permitemnos assim afirmar que os espectros polínicos da margem Ibérica reflectem
uma imagem integral da vegetação que coloniza o continente adjacente.
Este estudo permitiu ainda demonstrar que os espectros polínicos
marinhos
discriminam
claramente
ambas
as
comunidades
vegetais
mediterrânicas e atlânticas que colonizam as zonas situadas a sul e a norte
da Peninsula Ibérica, respectivamente.
Para além disso, este estudo testemunha, tal como noutras regiões do
mundo, que os grãos de pólen de Pinus encontram-se sobrerepresentados
nos sedimentos marinhos, e que devem por isso ser retirados no somatório de
base o qual é geralmente utilizado no cálculo das percentagens polínicas.
Este trabalho permitiu ainda determinar os tipos de mecanismos
associados à transferência dos grãos de pólen do continente para a
margem Ibérica.
Actualmente esta margem é dominada por ventos vindos de noroeste
os quais impedem o transporte aéreo dos grãos de pólen do continente
para o oceano. Por outro lado, a margem Ibérica localiza-se nas
proximidades de uma série de bacias hidrográficas importantes tais como o
Tejo, o Sado, o Douro e o Minho, as quais facilitam o transporte fluvial destes
grãos para o mar aberto.
A comparação dos resultados obtidos no estudo das concentrações
polínicas totais das amostras de superfície com modelos conceptuais
relacionados com a dinâmica sedimentar das partículas finas ao longo desta
margem permitiu-nos propôr um padrão para a dispersão polínica nesta
região.
No noroeste da Península Ibérica, os grãos de pólen são incorporados
nas camadas nefelóides e transportados para o mar pelos rios Douro e
Minho. Ao chegar à plataforma continental a maioria dos grãos são
bloqueados pelos afloramentos rochosos ai existentes. No inverno, durante
condições de "downwelling”, os grãos de pólen serão transportados para
norte, depositando-se primeiramente no complexo silto argiloso do Douro, no
complexo silto argiloso do Minho, e finalmente para o mar aberto (oeste).
Apenas uma parte ínfima destes grãos é transportada directamente do rio
para o mar aberto, principalmente durante períodos de fortes tempestades.
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F. Naughton, 2007
Durante o verão, as condições de "upwelling” impedem o transporte de
grãos de pólen para o mar aberto, pelo que apenas uma pequena
quantidade poderá ser transferida através de filamentos transversais.
No sul da margem Ibérica, os grãos de pólen são libertados pelos rios
Tejo e Sado e são parcialmente depositados na plataforma continental. Estes
grãos são posteriormente transportados para sul e em direcção do mar
aberto
por
correntes
litorais
e
oceânicas
durante
condições
de
“downwelling”, ou ainda através dos canhões submarinos.
Para além da credibilidade das assinaturas polínicas marinhas actuais,
este estudo demonstrou ainda que as sequências polínicas marinhas da
margem
oeste
Ibérica
constituem
uma
ferramenta
indispensável
à
correlação directa oceano-continente visto que este tipo de estudos se
baseia numa base cronológica comum. Desta forma, foi efectuada a
comparação entre os resultados obtidos pela análise polínica de alta
resolução temporal, efectuada numa sequência marinha composta (a qual
inclui as sondagens MD99-2331 e MD03-2697) com dados obtidos noutras
sondagens marinhas desta mesma margem e ainda com várias sequências
polínicas continentais. Esta comparação permitiu ainda documentar as
variações do coberto vegetal ao longo da Península Ibérica para os últimos
25 000 anos e mostrar que a sequência marinha Galega apresenta uma
assinatura polínica representante essencialmente das baixas e médias
altitudes do noroeste da Península Ibérica.
A correlação directa entre indicadores paleoclimáticos marinhos
(conteúdo
em
IRD,
percentagens
de
foraminíferos
planctónicos
característicos de ambientes polares e valores em δ18O dos foraminíferos
planctónicos) e continentais (pólens) revelou ainda que a vegetação do
noroeste da Península Ibérica respondeu contemporaneamente aos eventos
climáticos detectados no Atlântico Norte, nomeadamente aos eventos de
Heinrich 2 e 1 (H2 e H1) (26 000-24 380 anos cal BP e 15 900-18 500 anos cal
BP), ao último máximo glaciar (LGM) (24 300 e 18 500 anos cal BP), ao evento
quente Bölling Alleröd (B-A) (13 200-15 900 anos cal BP) e ao evento frio Dryas
recente (YD) (11 600-13 200 anos cal BP).
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F. Naughton, 2007
A resposta da vegetação à variabilidade climática que caracteriza os
eventos de Heinrich 2 e 1 é complexa e caracterizada essencialmente por
duas fases distintas principalmente nas baixas e médias altitudes do noroeste
da Península Ibérica.
O início de cada um dos eventos é caracterizado por uma forte
regressão da floresta de pinheiros (Pinus) e pela expansão de urze
(Ericaceae incluindo Calluna) sugerindo que esta fase inicial seria bastante
fria e húmida. A segunda fase é caracterizada pela re-expansão da floresta
de pinheiros, a qual indica um episódio ligeiramente mais quente do que o
anterior. Para além disso, o desenvolvimento de plantas semi-desérticas
sugere um aumento da aridez durante esta segunda fase. Neste trabalho foi
ainda possível demonstrar que o evento de H1 é o equivalente marinho do
episódio designado por "Oldest Dryas” no continente.
O último máximo glaciar (LGM) foi essencialmente dominado por uma
vegetação herbácea no noroeste da Península Ibérica. Este período é ainda
caracterizado pela expansão pinheiros (Pinus) e presença esporádica da
floresta decídua (Quercus, Corylus e Alnus). A presença esporádica de
árvores temperadas permite-nos inferir que não só o sul mas também a
região norte terá agido como uma zona refugio para essas árvores durante o
LGM. A presença de urze (Ericaceae incluindo Calluna) sugere que o
noroeste foi, tal como o sudoeste da Península Ibérica, dominada por
condições húmidas.
A nossa sequência Galega mostra ainda que a expansão de
carvalhos (Quercus deciduous) ocorreu de forma mais rápida e intensa nas
baixas e médias altitudes do que nas regiões altas do noroeste da Península
Ibérica durante o episódio quente que caracteriza o evento Bölling-Allerød.
Isto sugere que a floresta temperada dessas regiões baixas reagiu, tal como
na zona sul, de forma mais rápida à variabilidade climática do Atlântico
Norte durante esse período interestadial.
O Dryas recente foi caracterizado pelo aumento de espécies
pioneiras (compostas essencialmente por Betula) e pela expansão de
gramíneas e plantas semi-desérticas (Artemisia e Ephedra) em detrimento da
floresta de carvalhos, nas baixas e médias altitudes do noroeste da Península
296
F. Naughton, 2007
Ibérica. Contudo, este evento é ligeiramente menos acentuado nas baixas e
médias do que nas altas altitudes do noroeste Ibérico.
A resposta da vegetação decídua à melhoria climática que
caracteriza o início do Holocénico aparenta ser mais rápida nas baixas e
médias do que nas altas altitudes do noroeste da Península Ibérica. No
entanto, a sucessão de árvores (Juniperus, Betula, Pinus Quercus, Corylus e
Alnus) observada nas várias altitudes da região nortenha é semelhante e
bastante diferente daquela que caracteriza a região sul.
No entanto, a expansão de deciduous Quercus na região sul ocorreu
tal como nas regiões baixas e médias do noroeste da Península Ibérica de
forma mais rápida provavelmente como resposta à forte densidade de
zonas refugio durante o precedente LGM nestas zonas.
A correlação directa entre indicadores paleoclimáticos continentais e
marinhos permitiu-nos propôr dois eventuais mecanismos responsáveis pelo
padrão complexo deixado pelos eventos de Heinrich no noroeste da
margem Ibérica.
Este estudo permitiu constatar a presença de duas fases distintas na
vegetação do noroeste Ibérico as quais parecem estar intimamente ligadas
ao sinal complexo deixado pelos eventos de Heinrich (H4, H2 and H1) ao
longo da margem Ibérica.
A primeira fase é marcada por condições oceânicas de superfície
muito frias (evidenciada pelas associações de foraminíferos planctónicos,
δ18O e estimativa da temperatura baseada nas alcanonas) e pela virtual
ausência de icebergues no noroeste da margem Ibérica, assim como por um
arrefecimento continental extremo o qual é revelado pelo forte declínio da
floresta de Pinus. Esta primeira fase é ainda marcada pela expansão Calluna
e por um aumento da concentração polínica total sugerindo um aumento
da humidade na Península Ibérica.
A segunda fase, associada à chegada máxima dos icebergs a esta
margem, é caracterizada por condições oceânicas de superfície e
continentais (expansão de Pinus) menos frias e por um aumento da aridez
(representada pelo desenvolvimento de plantas semi-desérticas).
297
F. Naughton, 2007
O impacto do evento H3 no noroeste da Península Ibérica é peculiar
reflectindo condições húmidas durante quase todo o episódio. No entanto,
tal como acontece nos outros eventos de Heinrich (H4, H2 e H1), a
diminuição substancial da temperatura da massa de água de superfície e
da atmosfera (regressão da floresta de Pinus) precede a chegada máxima
dos icebergs durante este evento atípico.
Ao longo deste trabalho foram propostos dois mecanismos principais
para explicar este padrão complexo. O drástico arrefecimento atmosférico
(declínio da floresta de Pinus) detectado durante os eventos H4, H3, H2 e H1
resulta provavelmente da introdução de grandes quantidades água doce
provenientes da fusão dos icebergues no Atlântico Norte a qual provocou a
interrupção da circulação termohalina do Atlântico Norte (MOC). Esta
interrupção da MOC foi seguida por rápidas reorganizações entre o oceano
e a atmosfera as quais favoreceram a transmissão das condições frias para o
noroeste da Península Ibérica.
Sobreimpostas
a
este
mecanismo
oceanográfico,
variações
semelhantes aos modos negativo e positivo do índice da Oscilação Norte
Atlântica (NAO-like) parecem ter tido um papel crucial no padrão complexo
deixado pelos eventos típicos de Heinrich no noroeste Ibérico. Durante a
primeira fase, a predominância do modo negativo da NAO-like explicaria o
aumento na precipitação invernal e importantes descargas fluviais nesta
região, as quais favoreceriam o transporte de grãos de pólen do continente
para o oceano. Estas condições provocariam ainda o aumento da
temperatura superficial do Atlântico Norte a latitudes situadas acima dos 45°50°N, facilitando a fusão dos icebergues provenientes da calote glaciar da
“Laurentide” na cintura de IRD (IRD belt) impedindo o seu transporte para as
latitudes médias do Atlântico Norte. A fusão dos icebergues produziria ainda
um arrefecimento da massa de água superficial nessa zona. A dominância
do modo negativo da NAO-like produziria ainda uma diminuição da
temperatura da camada de água superficial numa zona estreita próxima da
costa entre o noroeste da margem Ibérica e o Grande Mar do Norte
(Greater North Sea) favorecendo a amplificação das condições frias na
nossa zona de estudo. Durante a segunda fase dos eventos de Heinrich
típicos (H4, H2 e H1), a dominância do modo positivo da NAO-like poderiam
298
F. Naughton, 2007
ter provocado a intensificação e a migração para norte dos ventos de oeste
provocando um aumento da aridez na Europa incluindo a zona da Península
Ibérica. Estas condições produziram condições oceânicas de superfície
relativamente quentes nas médias latitudes do Atlântico Norte (~ 20°- 40°N)
favorecendo a migração dos icebergues para sul e a sua fusão nessa região
(incluindo a margem Ibérica). Apesar dos icebergues atingirem a margem
Ibérica,
as
condições
atmosféricas
prevalecentes
(NAO-like
positivo)
favoreceriam um aquecimento da massa de água superficial entre o
noroeste da margem Ibérica e o Grande Mar do Norte latitudes médias do
Atlântico Norte mascarando por isso o arrefecimento causado pela fusão
dos icebergues na zona de estudo.
A predominância de condições húmidas durante o evento atípico,
H3, poderia ser explicada pela permanência de ventos fracos vindos de
oeste, nesta região.
A comparação directa entre condições no continente e no oceano
revelou ainda que durante o LGM, apesar das condições oceânicas de
superfície serem quentes, a expansão da floresta temperada foi largamente
reduzida
quando
comparada
com
os
episódios
interestadiais
que
caracterizam o MIS3 tardio (MIS3-Marine isotopic Stage 3).
Neste trabalho são propostos três mecanismos de forma a explicar
esta disparidade entre as condições oceânicas e atmosféricas tais como a)
o aumento do albedo o qual teria provocado uma amplificação do
arrefecimento produzido pela baixa insolação de verão no Hemisfério Norte,
b) o forte contraste sazonal das altas latitudes do Atlântico Norte como
resultado de um arrefecimento substancial da temperatura de inverno, e c)
a diminuição da concentração de CO2 na atmosfera. No entanto, a
humidade relativa que dominou o oeste da Península Ibérica durante o LGM
parece resultar do aumento da intensidade da MOC em relação aos
eventos precedente e antecedente a este, ou seja, o H2 e o H1.
Para além do último período glaciário interessámo-nos ainda em
compreender a variabilidade climática milenar que ocorreu durante a última
deglaciação (19 500-7000 anos cal BP) nas médias latitudes do Atlântico
299
F. Naughton, 2007
Norte e no noroeste da Península Ibérica. Este período é caracterizado por
um aumento da insolação de verão no Hemisfério Norte e pela forte
redução do volume de gelo acumulado nos pólos. A deglaciação engloba
o final do LGM, o evento H1, o interestadial Bölling-Alleröd (B-A), o Dryas
recente (YD) e inicio do Holocénico (11 500-8 000 anos cal BP).
O aquecimento continental (expansão de Quercus deciduous) e
oceânico que caracteriza o evento Bölling-Alleröd (B-A) foi favorecido pelo
aumento da insolação de verão das latitudes médias do Hemisfério Norte e
pela intensificação da MOC. A expansão máxima de Quercus deciduous, a
qual reflecte um aquecimento continental extremo ocorreu há cerca de
14 000 anos cal BP e é síncrona do pico máximo da temperatura na
Gronelândia e do episódio de subida súbita do nível do mar (MWP 1A). Este
aquecimento severo do Hemisfério Norte poderá ter sido o impulsionador
deste drástico evento designado “Meltwater Pulse 1A”.
A subsequente diminuição da floresta de Quercus no noroeste da
Península Ibérica contemporânea da diminuição da temperatura da massa
de água superficial no oceano adjacente caracteriza o Dryas recente (YD)
nessa região. A redução em vez da interrupção da MOC e o aumento da
insolação de verão das latitudes médias do Hemisfério Norte favoreceram a
redução em vez do total declínio da floresta decídua de Quercus no
noroeste Ibérico. Para além do arrefecimento, a expansão de plantas semidesérticas reflecte o aumento da aridez continental. Este incremento das
condições áridas parece resultar da dominância do modo positivo da NAOlike na região Norte Atlântica.
O máximo térmico do Holocénico (HTM), caracterizado por uma forte
insolação de verão das latitudes médias do Hemisfério Norte, foi detectado
no noroeste da Península Ibérica entre 11 700 e 8 200 anos cal BP e é
contemporâneo da expansão máxima da floresta de Quercus deciduous
durante o actual interglaciário.
Por volta dos 8 200 anos cal BP, o rápido arrefecimento da massa
oceânica de superfície das médias latitudes do Atlântico Norte e a
diminuição da floresta decídua de Quercus e Corylus no noroeste da
Península Ibérica, reflectem o evento frio “8.2 ky” nessa região.
300
F. Naughton, 2007
Este episódio resulta da culminação de uma série de episódios
relacionados com o colapso da calote glaciar da “Laurentide” a qual
provocou a amplificação da diminuição da temperatura na Europa e na
Gronelândia.
Após o evento “8.2 ky” a diminuição gradual da floresta temperada é
contemporânea da diminuição da temperatura induzida pela diminuição
da insolação de verão das latitudes médias do Hemisfério Norte. Isto sugere
que a regressão da floresta temperada parece ter sido mais afectada pelas
variações orbitais do que pelo impacto antrópico.
A baixa resolução das sondagens marinhas profundas para o período
Holocénico impediu a detecção da variabilidade climática milenar durante
o actual interglaciário nas médias latitudes do Atlântico Norte, e em
particular de identificar o evento multi-secular que ocorreu entre ~8 600 e 8
000 anos cal BP. De forma a resolver esta lacuna foi efectuada a análise
polínica de uma sondagem marinha pouco profunda com cerca de 2.72 m
a qual foi recolhida na plataforma continental do noroeste de França (VK0358Bis). As variações do coberto vegetal e a estimativa dos parâmetros
climáticos
obtidos
nesta
sondagem,
permitiram
detectar
variações
climáticas de escala orbital e sub-orbital, para os últimos 8 850 anos, nesta
região.
O declínio gradual da floresta temperada juntamente com a
diminuição da temperatura de verão (MTWA-mean temperature of the
warmest month) desde 8 855 até, pelo menos, 2 000 anos cal BP é
contemporâneo
do
arrefecimento
progressivo
da
temperatura
na
Gronelândia, assim como, da redução dos valores de insolação de verão
nas latitudes médias do Hemisfério Norte. Ao mesmo tempo, a diminuição da
sazonalidade segue, tal como seria de esperar, o aumento da precessão.
Entre 8 739 e 8 387 anos cal BP, a floresta de Corylus expande-se em
detrimento do Quercus deciduous, como resposta a uma amplificação do
contraste sazonal, resultante da expansão de gelo marinho invernal nas altas
latitudes do Atlântico Norte, contrariando o padrão geral de forçagem
orbital. Este forte contraste sazonal, resulta da expulsão drástica de água
301
F. Naughton, 2007
doce dos lagos de “Agassiz” e “Ojibway” e, da gradual redução da MOC. O
súbito declínio da floresta de Corylus, entre 8 387 e 8 062 anos cal BP, marca
tal como em outras regiões da Europa, o evento frio “8.2 kyr”, no noroeste de
França. Este episódio, foi provavelmente produzido pela redução severa da
MOC a qual está interligada com um fluxo mínimo da “Iceland-Scotland
Overflow Water” (ISOW) provocando uma diminuição suplementar da
temperatura invernal, na Europa e na Gronelândia. No entanto, apesar do
forte declínio da floresta de Corylus o contraste sazonal permaneceu
elevado durante este evento. De facto, o forte contraste sazonal registado
entre 8 739 e 8 062 anos cal BP reflecte o evento multi-secular "8.6-8.0 kyr”
nesta região.
Após os estádios finais de expulsão de água dos lagos “Agassiz” e
“Ojibway”, o clima torna-se relativamente mais estável. No entanto, estão
registados uma série de ligeiros e rápidos episódios frios os quais, estão
associados a um pequeno arrefecimento invernal e a um ligeiro aumento da
precipitação.
Para além da análise polínica, foi ainda efectuado um estudo das
associações dinocistos e de gastrópodes do tipo Turritella communis o qual
permitiu detectar variações regionais tais com: a) migração para sul da zona
biogeográfica marinha Boreal, entre 8 739 e 8 479 anos cal BP, como
resposta ao aumento da área coberta por gelo marinho no Atlântico Norte
permitindo o estabelecimento das comunidades bentónicas (T. Communis)
no noroeste da margem françesa e, b) a abertura do canal da mancha
entre 8479 e 8387 anos cal BP a qual provocou grandes modificações
ambientais nesta região e consequente mortalidade da T. communis assim
como
a
forte
diminuição
das
percentagens
do
Lingulodinium
machaerophorum.
A variabilidade climática tem uma forte influência na evolução dos
sistemas costeiros principalmente desde a última deglaciação. De facto, o
aumento da insolação de verão das altas latitudes do Hemisfério Norte
provocou a diminuição do volume de gelo, e como consequência um
302
F. Naughton, 2007
aumento do nível do mar global o qual teve um forte impacto na evolução
dos sistemas costeiros mundiais.
Na tentativa de compreender o impacto destes mecanismos forçadores
globais, nomeadamente o clima e o nível do mar na evolução
geomorfológica dos sistemas costeiros do noroeste da Península Ibérica, foi
efectuado um estudo multidisciplinar em duas sondagens recolhidas na
embocadura do estuário do Douro (noroeste de Portugal).
A primeira metade do período Holocénico (10 720 e 6 530 anos cal BP)
é caracterizada por uma floresta de Pinus-Quercus-Alnus reflectindo um
clima quente e húmido na bacia do Douro. Durante este período, o
aumento gradual de foraminíferos bentónicos característicos de zonas de
plataforma e talude continental, assim como a presença de equinodermes
sugere o aumento gradual do nível do mar nessa zona. Os grãos de pólen,
representantes da vegetação regional que colonizava a bacia hidrográfica,
foram transportados pelo rio ao longo do antigo canal principal e
depositados na zona de estudo a qual seria nesse momento uma antiga
margem do rio.
Há cerca de 6 530 anos cal BP uma drástica variação ambiental,
associada ao desaceleramento da subida do nível do mar e aumento do
hidrodinamismo do rio, é marcada pela deposição de cerca de 5 m de
cascalho na parte sul deste estuário. O estabelecimento desta barreira
cascalhenta na parte sul deste estuário ocorreu entre 6 530 e 1 500 anos cal
BP e provocou a migração para norte do canal principal impedindo o
transporte e a deposição dos grãos de pólen representantes da vegetação
regional na zona de estudo após 1 500 anos cal BP.
A formação desta barreira e a consequente migração para norte do
canal principal do rio contribuíram para o alargamento da embocadura
deste estuário.
303
F. Naughton, 2007
304
F. Naughton, 2007
Anexos
Anexo A
Climate variability of the last five isotopic interglacials:
direct land-sea-ice correlation from the multiproxy analysis of north
western Iberian margin deep-sea cores.
S. Desprat, M.F., Sánchez Goñi, F., Naughton, J.-L., Turon, J. Duprat, B. Malaizé, E. Cortijo and
J.-P. Peypouquet
In press in The climate of past interglacials. Elsevier publications.
Climate variability of the last five isotopic interglacials:
direct land-sea-ice correlation from the multiproxy analysis of north western
Iberian margin deep-sea cores.
S. Desprata*, M.F., Sánchez Goñia, F., Naughtonb, J.-L., Turonb, J. Dupratb, B. Malaizéb, E. Cortijoc
and J.-P. Peypouqueta
a
Ecole Pratique des Hautes Etudes, Paléoclimatologie et Paléoenvironnements marins, Département
de Géologie et Océanographie, Université Bordeaux 1, Avenue des Facultés, 33405 Talence, France
b
Département de Géologie et Océanographie, Université Bordeaux 1, Avenue des Facultés, 33405
Talence, France
c
Laboratoire des Sciences du Climat et de l’Environnement, LSCE-Vallée, Bât. 12, avenue de la
Terrasse, F-91198 Gif-Sur-Yvette cedex, France
*corresponding author: [email protected]
Abstract
The last five marine isotopic interglacials (Marine Isotope Stages 11, 9, 7, 5 and 1) were investigated
in Iberian margin deep-sea cores, using terrestrial (pollen) and marine (planktic foraminifera, benthic
and planktic oxygen isotopes) climatic indicators. This work shows that the climatic variability
detected on the continent is contemporaneously recorded in the ocean, but temperature changes are not
in phase with ice volume variations. The comparison of the different marine isotope stages highlights a
common pattern of climatic dynamic within these interglacials. This dynamic is characterized by three
major climatic cycles, related to orbital cyclicity, on which suborbital climatic fluctuations are
superimposed. Particularly, suborbital events interrupt the deglacial warming associated to
Terminations IV to I and the second major warm period of each isotopic interglacial as well as the
transitions towards glacial marine isotope stages. MIS 7 displays a short first warm period (~8 ka)
followed by a striking cold and dry period succeeded by a new strong warmth. In contrast, MIS 11
presents the longest (~31 ka) period of the last 450,000 years.
1. Introduction
Forecasting the future climatic evolution of the current interglacial period is a great challenge.
Before that, it is necessary to determine the evolution of the past interglacials and evaluate the
response of different components of the Earth’s climatic system. Due to the Earth’s astronomical
configuration, Marine Isotope Stage (MIS) 11 is the best candidate to be the analogue of MIS 1.
However, characterising the climatic evolution over different situations of insolation forcing will allow
us a better understanding of climate dynamics during interglacial periods. The continental
paleoclimatic records covering the last 425,000 years are rare and often fragmentary, and their
chronologies are difficult to establish. This impedes the comparison of the climatic changes detected
on land with those identified in the oceanic and cryospheric realms.
We present, here, the first direct land-sea-ice correlation for the last five isotopic interglacials
(MIS 11, 9, 7, 5 and 1). The main purpose of this work is to document the climatic variability of these
periods and to assess the phase relationship between the responses of the different Earth’s
environments –continent, ocean and ice– to climatic changes in order to discern analogies and
differences between them. For that, a multiproxy study (pollen, assemblages of planktic foraminifera,
and planktic and benthic δ18O) was performed from several NW Iberian margin deep-sea cores. By
comparing the last five isotopic interglacials, we will highlight, on the one hand, the similarities of
their climatic dynamic despite different astronomical forcing and, on the other hand, the dissimilarities
concerning duration, warmth magnitude and forest succession of the warm periods.
2. Present-day environmental setting and pollen signal in the Iberian margin
The Iberian margin deep-sea cores were retrieved ~100 km off the Galician coast at ~2,100 m
of water depth (Fig. 1). This site is at present under the influence of the North Atlantic Deep Water.
The north western Iberian climate is considered temperate and humid as a result of the influence of
dominant Atlantic winds over the year. Mean annual temperature is 12.5°C (Mean Temperature of the
Coldest Month, MTCO = 5-12°C; Mean Temperature of the Warmest Month, MTWA = 17-22°C) and
mean annual precipitation is between 1000 and 2000 mm.an-1 (Atlas Nacional de España, 1992). This
region, incised in the north by the Rias Baixas valleys (Galician coast basin) and the Miño-Sil river
(Sil basin) (Atlas Nacional de España, 1993) and crossed in the south by the Douro river, belongs to
the Eurosiberian and sub-Mediterranean regions (Ozenda, 1982). At present, deciduous oak woodlands
(Quercus robur, Q. pyrenaica and Q. petraea), heaths (Ericaceae including Calluna), brooms
(Genista) and gorses (Ulex) dominate the vegetal cover of north western Iberia (Alcara Ariza et al.,
1987).
Studies on present-day pollen deposition in marine sediments show that pollen grains reach
marine sites from the adjacent continent by both fluvial and aeolian transport and subsequent sinking
through the water column (Chmura et al., 1999; Dupont and Wyputta, 2003; Heusser, 1978; Heusser
and Balsam, 1977). Further, it is suggested that cores located near continental regions with well
developed hydrographic basins and prevailing offshore winds, as it is the case of our cores, mainly
recruit pollen from rivers (Dupont and Wyputta, 2003; Heusser, 1978; Turon, 1984). The north
western Iberian rivers, mainly the Douro and Miño and slightly the Rias Baixas, provide sediments to
the shelf area. On the shelf, the fine-particle sediments in suspension are transported northwards by
poleward currents, and some are deposited in the Douro and Galicia Mud Patches (Dias et al., 2002).
Extreme storm events can produce re-suspension of some sediment from the mud patches and transport
of sediments off the shelf can occur (Jouanneau et al., 2002; Vitorino, 2002). Pollen grains belonging
to the fine-particle fraction have a similar behaviour as fine sediments during the sedimentary
processes (Chmura and Eisma, 1995; Muller, 1959). This suggests that pollen grains preserved in our
Iberian margin cores mainly come from the Galician and Douro fluvial basins. The comparison of
marine and continental modern pollen samples with the present-day Iberian vegetation shows that our
marine pollen records represent an integrated image of the regional vegetation of the north western
part of the peninsula (Naughton et al., in prep).
3. Material and method
Records of pollen and classical climate indicators (planktic foraminafera assemblages, benthic
and planktic δ18O) for the last five isotopic interglacials derive from three north western Iberian margin
deep-sea cores (MD01-2447, MD03-2697, MD99-2331) (Fig. 1). They were retrieved at the same
coordinates on board of the Marion Dufresne oceanographic ship, using the giant corer CALYPSO. As
shown in figure 2, the intervals corresponding to MIS 11, 9 and 7 were studied in core MD01-2447.
The beginning of stage 9, being unfortunately disturbed in this core, was studied in the twin core
MD03-2697. We present therefore a composite record of MIS 9 built from the correlation of several
marine proxies analysed in the twin cores (lightness L*, CaCO3 content, percentages of N. pachyderma
s., coiling ratio of Globorotalia truncatulinoides) (Desprat, 2005a). MIS 1 and 5 records come from
the third core MD99-2331 (Naughton et al., in prep; Sanchez Goñi et al., 2005).
3.1 Pollen analysis
Each interval corresponding to MIS 11, 9, 7, 5 and 1 was subsampled for pollen analysis at 10
or 5 cm intervals. The sample preparation technique followed the procedure described by de Vernal et
al. (1996) improved at the "Département de Géologie et Océanographie", University Bordeaux I
(Desprat, 2005a). After chemical and physical treatments (cold HCl, cold HF and sieving through 10
µm nylon mesh screens), the residue was mounted unstained in glycerol. Pollen was counted using a
Zeiss Axioscope light microscope at 500 and 1250 (oil immersion) magnifications. A minimum of 100
pollen grains excluding the over-represented Pinus grains in marine deposits (Heusser and Balsam,
1977; Turon, 1984), were counted in each of the 327 samples analysed. The pollen percentages for
each taxon are based on a main pollen sum that excludes Pinus, aquatic plants, pteridophyte spores and
indeterminable pollen. Pinus percentages were calculated from the main sum plus
Pinus. Spores and aquatic pollen percentages were obtained from the total sum (pollen + spores +
indeterminables + unknowns).
3.3 Isotopic analyses
The sampling resolution interval oscillates between 20 and 2 cm for Cibicides wuellerstorfi and
Melonis barleeanus benthic foraminifera and Globigerina bulloides planktic foraminifera. Each
specimen has been picked up within the 250 - 315 µm grain-size fraction, and cleaned with distilled
water. The preparation of each aliquot (4 to 8 specimens, representing a mean weight of 80 µg) has
been done using the Micromass Multiprep autosampler, using an individual acid attack for each
sample. The CO2 gas extracted has been analyzed against NBS 19 standard, taken as an international
reference standard. The isotopic analyses of core MD01-2447 and MD99-2331 have been carried out
at the "Département de Géologie et Océanographie" (Bordeaux 1 University, France), using an Optima
Micromass mass spectrometer, and those of core MD03-2697 were performed at the "Laboratoire des
Sciences du Climat et l’Environnement" (Gif-sur-Yvette, France), using a delta plus Finnigan isotope
mass spectrometer. All the isotopic results are presented versus PDB. The mean external
reproducibility of powdered carbonate standards is ±0.05‰ for oxygen.
The δ18O values for Cibicides wuellerstorfi and Melonis barleeanus were adjusted by +0.64 per
mil and +0.36 per mil, respectively, to account for species-dependent departure from isotopic
equilibrium (Duplessy et al., 1984; Graham et al., 1981; Jansen et al., 1988; Shackleton and Opdyke,
1973).
3.4. Chronological framework
The age model of the intervals corresponding to MIS 11 and 9 is based on the graphical
correlation of the benthic δ18O curve with the Low Latitude Stack of Bassinot et al. (1994) (Desprat,
2005a; Desprat et al., 2005b). The chronology of MIS 7 section also derived from a graphical
correlation but in this case using the benthic-stack of Martinson et al. (1987) (Desprat et al.,
submitted).
The chronologies of MIS 5 and 1 are, in contrast, independent of the astronomical calibration.
That of the last isotopic interglacial is based on the correlation of the major climatic phases detected in
core MD99-2331 with those identified and dated in the southern Iberian margin core MD95-2042
using the MD95-2042 chronology of Shackleton et al. (2002) (Sánchez Goñi et al., 2005). For the
interval corresponding to the last 25,000 years, the age model was established using the chronology of
the climatic episodes identified in other North Atlantic records and the ages assigned to several welldated botanical events in the Iberian Peninsula (Naughton et al., in prep).
4. The climatic variability of the last five isotopic interglacials in and off NW Iberia
4.1. General climatic dynamic
During the last five isotopic interglacials, the warm periods in north western Iberia are
characterized by the development of the temperate and humid forest, principally deciduous oak. In
turn, open vegetation dominated by Poaceae and Asteraceae, with some semi-desert plants, or mainly
composed by heathland, expands during cold periods. The recorded vegetation changes indicate that
climate has strongly oscillated during the previous isotopic interglacials.
The climatic evolution detected on the continent parallels the oceanic changes reflected by
marine proxies (planktic foraminifera δ18O and percentages). Indeed, each cold episode is marked by
an increase of the percentages of the polar planktic foraminifera N. pachyderma s. and heavier planktic
δ18O values. The record of planktic foraminifera assemblages is only available for MIS 11, 9 and 7. It
shows that the tropical and summer subtropical species generally reach their maximal development
during the warm periods detected on the continent.
During each of the Terminations I to IV, an abrupt cold event interrupts the development of the
temperate and humid forest associated to the deglacial warming: Younger Dryas, post-Zeifen stadial,
MD47-7-S1 and MD47-9-S1 (Fig. 3). These cold events have a clear imprint in the oceanic realm, in
particular that of Termination IV which is slightly marked on the continent. These coolings appear of
different magnitude and during Termination II, III and IV, they occur at the onset of minimum ice
volume. For example, MD47-8-I1/MD47-7-S1 cycle shows the strongest amplitude of vegetation
changes. The occurrence of such an episode during Termination V is still unknown because our
sedimentary core does not cover the whole MIS12/MIS11 transition.
As suggested by the long European pollen sequences (Reille et al., 2000; Reille et al., 1998;
Tzedakis et al., 2001), our direct land-sea-ice correlation confirms that each isotopic interglacial is
characterized by three major warm periods on the continent associated to low ice volume, in response
to the astronomical forcing (Table 1, Fig. 3). Indeed, the major forested periods in north western Iberia
are associated to low ice volume contrasting with the open vegetation phases related with ice cap
development. Nevertheless, our direct land-sea-ice correlation puts forward that the ice volume
changes are not synchronous with the temperature shifts on the continent and in the ocean. As
observed by Sánchez Goñi et al. (1999, 2005), Shackleton et al., 2002 and Tzedakis et al. (2004), the
limits of isotopic substages do not correspond to those of the climatic phases detected in western
Iberia. For example, the Eemian in Iberia does not correspond to the entire MIS 5e (Sánchez Goñi et
al., 1999, 2005; Shackleton et al., 2002).
During the isotopic interglacials MIS 11, 9, 7 and 5, the first major warm periods (Vigo,
Pontedevedra, Arousa and Eemian) are marked by a more developed forest than the following ones.
This is particularly true for MIS 11. Indeed, the climate optimum of each stage, detected by the
maximal expansion of oak forest, Mediterranean plants and the maximal contraction of pine
woodlands, occurs during these first major warm periods. These climate optima are contemporaneous
to the ice volume minimum of each stage. However, MIS 7 presents another particularity likely related
to especially strong insolation maximum: the second major warm period (Ribeira) is also marked by
strong expansion of the temperate and humid forest, development of warm planktic foraminifera and
important ice volume decrease. This implies that Ribeira would be, at least, as warm as the Arousa
interglacial. Therefore, MIS 7 displays two climatic optima both associated to low ice volume as
shown by the benthic isotopic record.
The Vigo, Pontedevedra and Eemian interglacials are followed by a strong cold period (MD4711-S1, MD97-9-S2 and Mélisey I). In contrast, MIS 7 includes a suborbital cycle (MD47-7-S2/MD477-I2) between the Arousa interglacial and the strong cold period MD47-7-S3.
MD47-7-S3 was the coldest and driest period on the continent of the last five isotopic
interglacials, as indicated by the highest values of grassland and semi-desert taxa of our pollen record.
This phase is also marked by the strongest decrease of sea surface temperatures as shown by the
maximum percentages of N. pachyderma left coiling. This cold episode is also coeval with the largest
ice volume increase (MIS 7.4) of the last past isotopic interglacials, being similar to the glacial
maximum of MIS 8. These unusual very cold conditions and huge ice-sheet enlargement within an
isotopic interglacial, are also recorded by northern North Atlantic ODP sites 983 and 980 (McManus et
al., 1999; Channell et al., 1997). It is remarkable that this episode occurs during the most important
insolation minimum of the last 450,000 years.
Other suborbital events are superimposed to this orbital climatic variability. The Vigo
interglacial is marked by two cool events circa 417 and 402 ka, having a clear imprint on vegetation
and planktic foraminifera. More especially, the temperate and humid forest developments associated to
the second major warm periods of MIS 11, 9 and 5 (Moana, Sanxenxo, St-Germain I) are all
interrupted by a cold episode (MD47-11-S1, MD97-9-S2 and Montaigu, respectively). After our age
model, these cold events are short, between 2 and 4 kyrs. Moreover, during these episodes, the benthic
δ18O values become heavier, in particular during MD97-9-S2. This indicates a significant increase of
ice volume during the cold event within MIS 9c, as shown by Tzedakis et al. (2004). These cold
fluctuations are also clearly recorded in the Velay sequence (Reille et al., 2000). In contrast, MIS 7
does not show such a climatic event.
During the interglacial-glacial transitions MIS 7-MIS 6 and MIS 5-MIS 4, another climatic
cycle of minor order is detected. The cooling associated to the MIS 9-MIS 8 transition seems also to be
interrupted by a warm oscillation which needs to be confirmed by supplementary analysis.
Nonetheless, all these warm events are clearly recorded in the oceanic realm by light values of planktic
δ18O and an increase of warm planktic percentages.
In sum, in spite of the different astronomical configuration of the last five isotopic interglacials
a common climatic evolution pattern emerges. However, each warm phase is characterized by different
duration, amplitude, and forest succession.
4.2. Warmth amplitudes of the last 450,000 years
At present, discussions are open on the different amplitudes of warmth during the last 450,000
years. The results are often contradictory, depending on the regions and proxies concerned. Some
works suggest that the warmest phase of MIS 11 shows the highest temperatures of the last 500,000
years (Howard, 1997; Droxler and Farrel, 2000; Berstad et al., 2002). However, this idea is challenged
by many works (Hodell et al., 2000; Bauch et al., 2000; Kunz-Pirrung et al., 2002; McManus et al.,
2003) which have shown that MIS 11 was not warmer than today. The deuterium signal of Vostok ice
core records the highest temperatures for MIS 9 (Petit et al., 1999) but the recent results of EPICADome C ice core do not confirm this idea (EPICA community members, 2004). These new data also
show higher temperatures in Antarctica during MIS 11 than during the Holocene (EPICA community
members, 2004).
As previously shown, in our record, each isotopic stage displays its climate optimum during the
first warm period, excepting MIS 7 which presents a second optimum during Ribeira. The first
optimum of this isotopic interglacial (Arousa) is marked by the highest percentages of temperate and
humid trees. However, the Mediterranean plants are scantily represented during this interval and only
by evergreen Quercus which does not reveal clear Mediterranean conditions. In contrast, although the
Vigo, Ribeira, Eemian and Holocene interglacials are characterized by lower percentages of temperate
and humid forest than those during the Arousa interglacial, they record true Mediterranean species
such as Pistacia, Olea or Cistus. For this reason, it remains difficult to determine what period is the
warmest of the last five climatic cycles in northern Iberia.
The warm planktic foraminifera record of MIS 11, 9 and 7 indicates the highest sea surface
temperatures during MIS 11 climatic optimum. However, the development of these warm foraminifera
is only slightly stronger than during the other climatic optima. Therefore, the Vigo interglacial may be
the warmest period of the last 450,000 years but the difference of temperature does not appear large.
Moreover, the benthic isotopic signal does not display weaker ice volume during Vigo interglacial than
during Pontedevedra, Ribeira, Eemian or Holocene interglacials. The genesis of such a warm
interglacial during MIS 11 remains still a mystery in palaeoclimatology taking into account the weak
insolation forcing.
4.2. Duration of the forest phases
MIS 11 is marked by a long first major warm period, lasting ~31 ka after our age model. The
Vigo interglacial appears two times longer than the Eemian (~16 ka), and at least three times longer
than the Pontedevedra and Arousa interglacials (~11 and 8 ka, respectively). The Holocene began
10,000 years ago and it is already longer than the Arousa interglacial.
On the basis of the pollen analysis of the south western Iberian margin deep-sea core MD012443, Tzedakis et al. (2004) and Roucoux et al. (this volume) suggest a very short forest phase during
MIS 9e, lasting 3,600 years, after which Ericaceae expand. To bypass the difference in duration
inferred from the age models, we have tuned our planktic δ18O record to that of MD01-2443. After this
exercise, the resulting duration of the period Pontedevedra is of 13,000 years. It is possible that the
cold/arid event responsible for the shortening of the first warm period of MIS 9 in south western Iberia
is not shown by our sequence due to a too low resolution analysis. However, our results do not suggest
that this abrupt climatic deterioration detected in south western Iberia brings to an end the first MIS 9
forest period in the north western part of the Peninsula. In the same way, the Praclaux sequence
(Massif Central, France) also records a long warm period even if it includes a slight forest reduction
(Tzedakis et al., 2004).
Our observation confirms, despite the uncertainties associated to the age scale of our record,
that the warmest period of MIS 11 would be longer than those of the following isotopic interglacial
stages, and so far three times longer than the Holocene.
4.4. Forest successions
The warm periods of our record are marked by the development of the pioneer trees, principally
Betula, followed by the expansion of deciduous oaks and that of hazel trees. This is in agreement with
the classical vegetational succession during interglacial period in northern Europe, described by Van
der Hammen et al. (1971). This ideal succession sees at the latest stages the expansion of hornbeam,
beech, fir, finishing by the development of the boreal forest with spruce. In the north western Iberian
region, the boreal forest phase is never reached during the warm periods. Moreover, the expansion of
the latecomer trees is different from one period to another:
•
Abies strongly developed only at the end of the Vigo interglacial and St Germain Ic and it was only
present at the end of the Eemian and Bueu (the last forested period of MIS 9).
•
Carpinus betulus expanded in the second part of all first major warm phases when deciduous
Quercus decreased excepting during the Holocene. However, the hornbeam expansion was very strong
during the Eemian, smaller during the Pontedevedra and Arousa interglacials and very weak during the
Vigo interglacial. During the other warm periods, hornbeam had also its maximal expansion after that
of deciduous oak.
•
Fagus never had an important expansion at the end of the first major warm phases. Beech is only
sporadically recorded at the end of the Vigo and Pontedevedra interglacials. In contrast, during the
periods Bueu, Ribeira and Rianxo, Fagus plays an important role in the vegetal cover of the north
western Iberian region, always associated to Carpinus betulus. It is also noteworthy that beech
developed rapidly and strongly at the beginning of the Ribeira interglacial (Desprat et al., submitted).
The different behaviour of these three trees depending on the periods and regions has been
previously noticed by Tzedakis and Bennett (1995) and Tzedakis et al. (2001). Disentangling the
factors responsible for the settlement of tree species in given region and period is a difficult task. The
late expansion of some tree species can be linked: a) to their migration rate in relation with their own
dispersion mechanisms such as reproduction or seed scattering and with their competencies to develop
on more or less mature soils, b) to the distance with the glacial refugial zone, c) to the interspecific
competition, d) the individual response of each specie to a climatic change, and e) also to the direct
effect of the climate (Huntley and Webb, 1989; Huntley, 1996). The climate can also play an indirect
role in changing the interspecific relationships (Lischke et al., 2002). Small differences of climatic
conditions at the beginning of a warm phase can also influence the development of the late-expanding
trees (i.e. Carpinus, Abies, Fagus) (Tzedakis et al., 2001). Moreover, reduced diversity of taxa such as
Fagus, Carpinus and Abies may imply that they are more susceptible to disease or adverse climatic
conditions (Tzedakis et al., 2001). During a short warm period such as the Arousa interglacial, the
virtual absence of beech and fir in north western Iberia may be associated to a too short time for
migrating from faraway refugial areas to Iberian Peninsula, likely in relation with their own migration
mechanisms. Nevertheless, the biotic processes cannot explain the very late arrival of Fagus in north
western Iberia, approximately 25,000 years after the beginning of the Vigo interglacial, since it
developed only 7,000 years after the beginning of the Pontedevedra interglacial. Therefore, the
different timing and magnitude of the expansion of the late succession trees is somehow linked to the
inherent climatic conditions of each warm phase of the different isotopic interglacials, which are
related with the ice-sheet extension and the orbital parameters.
5. Conclusions
This work constitutes a new step in documenting the climatic variability of interglacial isotopic
stages. It provides the first direct land-sea-ice correlation of the last five isotopic interglacials (MIS 1,
5, 7, 9 and 11) from the multi-proxy analysis of three pollen-rich north western Iberian margin cores.
This record puts forward the phasing, previously identified during MIS 5, between changes in oceanic
surface conditions and continental climate during the previous isotopic interglacials. Despite the
differences of astronomical forcing, several similarities between these isotopic stages emerge: a) the
occurrence of three major climatic cycles, related to the orbital cyclicity, b) a climatic optimum during
the first major warm periods, associated to ice volume minimum, c) a suborbital cold event
interrupting the second major warm period, and d) a suborbital climatic instability during the glacialinterglacial and interglacial-glacial transitions.
The largest insolation oscillations controlling MIS 7 may explain the discrepancies between the
climatic variability of this isotopic interglacial and the observed general scheme: a second major warm
period at least as warm as the first one and preceded by a very cold and dry episode associated to an
unusual important ice volume. Another striking feature of this stage is the very short first warm period,
the Arousa interglacial, which is even shorter than the Holocene. In contrast, MIS 11 presents the
longest warm period (Vigo interglacial) of the last 450,000 years, three times longer than our present
interglacial.
Acknowledgements
Financial support was provided by IPEV and PNEDC French programs. We thank logistics and coring
teams on board of the R/V Marion Dufresne II during the Ginna, Geosciences and Picabia
oceanographic cruises and Marie-Hélène Castera, Karine Charlier, Olivier Ther and Françoise Vinçon
for invaluable technical assistance. This paper is Bordeaux 1 University, UMR-CNRS 5805 EPOC
Contribution n°.
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Tzedakis, P.C., Andrieu, V., de Beaulieu, J.-L., Birks, H.J.B., Crowhurst, S., Follieri, M.,
Hooghiemstra, H., Magri, D., Reille, M., Sadori, L., Shackleton, N.J., Wijmstra, T.A., 2001.
Establishing a terrestrial chronological framework as a basis for biostratigraphical comparisons.
Tzedakis, P.C., Roucoux, K.H., de Abreu, L., Shackleton, N.J., 2004. The duration of forest stages in
southern Europe and interglacial climate variability. Science 306, 2231-2235.
van der Hammen, T., Wijmstra, T.A., Zagwijn, W.H., 1971. The floral record of the Late Cenozoic of
Europe. In: Turekian, K.K. (Ed.), The Late Cenozoic glacial ages. Yale University press, New Haven,
pp. 391-424.
Vitorino, J., 2002. Winter dynamics on the northern Portuguese shelf. Part2: bottom boundary layers
and sediment dispersal. Progress in Oceanography 52, 155-170.
Figure 1: Location of the studied deep-sea cores MD01-2447, MD99-2331, MD03-2697.
Figure 2: Lightness record of the three Iberian margin deep-sea cores. Pollen analysis has been
performed in the intervals represented by grey areas. Hatched area corresponds to the disturbed
interval in core MD01-2447.
Figure 3: Direct correlation of continental and marine proxies from Iberian margin deep-sea cores.
From the left to the right: 1) Synthetic pollen diagram; 2) Percentages of warm planktic foraminifera
(only for MIS 7, 9 and 11) and Neogloboquadrina pachyderma left coiling; 3) Planktic δ18O curve; 4)
Benthic δ18O curves. For the last 25,000 years, the benthic isotopic data of core MD99-2331 being not
available, we present those of core MD01-2447. The correlation between both cores has been
performed using different marine proxies (lightness, CaCO3 content, abundance and coiling ratio of
Globorotalia hirsuta and G. truncatulinoides and percentages of N. pachyderma s.) (Desprat, 2005a);
5) June insolation at 65°N. Blue areas indicate cold periods.
Marine
Stages
Isotope
Isotopic events
Iberian major warm
periods
MIS 1
1.1
Holocene
MIS 5
5.1
St Germain II
5.3
St Germain I
5.5
Eemian
7.1
Rianxo
7.3
Ribeira
7.5
Arousa
9a
Bueu
9c
Sanxenxo
9e
Pontedevedra
11.1
Cangas
11.23
Moana
11.3
Vigo
MIS 7
MIS 9
MIS 11
Table 1: Major warm periods detected in north western Iberia during the last 425,000 years versus
marine isotopic stratigraphy. The time lags between the boundaries of isotopic substages and stages
and those of forested phases are not indicated in this table.
Anexo B
Paleoenvironmental evolution of estuarine systems during the
last 14000 years – the case of Douro estuary (NW Portugal)
T. Drago, M. C.Freitas, F.Rocha, J.Moreno, M.Cachão, F.Naughton, C.Fradique, F.Araújo,
T.Silveira, A.Oliveira , J.Cascalho, F.Fatela
In press in Journal of Coastal Research
Paleoenvironmental evolution of estuarine systems during the last 14000
years – the case of Douro estuary (NW Portugal)
T. Drago †, M. C.Freitas ‡, F.Rocha∞, J.Moreno‡, M.Cachão‡, F.Naughton†§, C.Fradique£, F.AraújoΦ,
T.Silveira†, A.Oliveira Ψ, J.Cascalho£, F.Fatela‡
† INIAP,
IPIMAR,
CRIPSUL
Olhão, 8700305, Portugal
tdrago@ipimar
.pt
‡ Dep. e Centro
Geol. Univ. Lisboa,
Lisboa, 1749-016,
Portugal
[email protected]
[email protected]
[email protected]
∞ Departamento
de Geociências,
Univ. Aveiro,
Aveiro, 3810193, Portugal,
[email protected]
§ DGO UMRCNRS 5808,
Univ. Bordeaux I,
Talence, 33405,
France,
f.naughton@epoc
.u-bordeaux1.fr
£ Museu Nacional
História Natural,
Lisboa, 1600-250,
Portugal
[email protected]
[email protected]
Φ Instituto
Tecnológico e
Nuclear,
Sacavém
2686-953
Portugal
[email protected]
Ψ Instituto
Hidrográfico,
Lisboa, 1100,
Portugal
anabela.oliveira
@hidrografico.pt
ABSTRACT
DRAGO, T.; FREITAS, M. C.; ROCHA, F. ; MORENO, J.; CACHÃO, M.; NAUGHTON, F.; FRADIQUE, C.;
ARAÚJO, F.; SILVEIRA, T.; OLIVEIRA, A.; CASCALHO, J.; FATELA, F., 2004. Paleoenvironmental
evolution of estuarine systems during the last 14000 years – The case of Douro estuary (NW Portugal)
(Proceedings of the 8th International Coastal Symposium), pg – pg. Itajaí, SC – Brazil, ISSN 0749-0208
The multidisciplinary study (sedimentology: texture, carbonates, organic matter, elemental composition,
mineralogy of sand and clay; micropaleontology: foraminifera, calcareous nannoplankton and pollen – and
radiocarbon dating) of three cores recovered from the Douro estuary infill has allowed to reconstruct the
paleoenvironmental guidelines of this area for the last 14000 years. The major environmental changes recognized
in the geological record of this area are principally associated with the Holocene transgression and the succession
of different climatic types. The transgressive sequence is expressed through several environmental changes,
related with a succession of different facies: a continental one till circa 9834 BP, followed by an alternated
marine/continental facies with an intensification of marine signature till 5750BP; particularly after 6050BP, the
environment is typically marine. After that and till present, it is followed by a regressive sequence (although
within a still positive eustatic trend).
ADITIONAL INDEX WORDS: environmental changes, estuarine systems, multi-proxies, sea-level rise
INTRODUCTION
Among depositional environments, estuarine systems are
privileged locals to study the continental and marine influences
that occurred in the recent past. The study of their sedimentary
record can be a useful tool in the understanding of the
environmental changes affecting the continental margin,
particularly since the last deglaciation and following sea level rise.
The present-day knowledge on the evolution pattern for the
Portuguese continental margin since the last millennia is based on
the work carried out by DIAS (1987) on the sedimentary dynamics
and recent evolution of the shelf and paleoenvironmental
reconstructions of southwestern coastal lagoons such as Santo
André and Melides (BAO et al., 1999; CRUCES et al., 1999;
FREITAS et al., 2002). The aims of this paper are thus, to present
the data resulting from multi-proxy studies of the Douro estuary
sedimentary record, in order to propose a general evolutionary
model for the last 14000 years and to provide additional
information on the paleoenvironmental evolution of the NW
Portuguese coast.
STUDY AREA
The Douro estuary is located in the northwestern Portuguese
coast, south of Oporto. It is elongated almost W-E, funnel-shaped
with 2,25km length, 1,25m maximum width and a mean depth of
about 5m. The estuary is partially sheltered from the ocean by a
sand spit, rooted in the southern margin, known as Cabedelo
(Fig.1). This is a common feature of all northern Portuguese
estuaries (Ave, Cávado, Lima, Minho), with the barrier differing
in length (in the case of Douro, it can attain 1km long) and
resulting from local reversal of littoral drift. Its morphology is
quite variable, dependant on the marine and fluvial hydrodynamic
conditions. In winter, during extreme floods, this barrier can be
overtopped, breached or even drowned for short periods that can
last some days or weeks.
Douro basin is located in the rainiest region of Portugal. It
presents a climate varying from humid to semi-arid; the humid
season occurs in October-March with 73% of total precipitation.
The annual mean discharge is 710m3/s, with a maximum of
3000m3/s and a minimum of 50 m3/s (LOUREIRO et al., 1986).
METHODS
This work is based on the study of three cores, which were
obtained by rotary drilling in the Douro estuary, one in the barrier
(core 2), known as Cabedelo and two in the back-barrier (cores 1
and 1B, 50 cm apart), inside de S. Paio Bay (Fig. 1).
Core 1 (+3.5/-4.4m height), 1B (-6.6/-16.2m height) and 2
(-15.16/-39.66m height) can be considered as representative of the
complete sedimentary sequence present in the southern margin of
Paleoenvironmental evolution Douro estuary (NW Portugal)
Douro estuary. Cores 1B and 2 reached basement, consisting of
weathered granite.
Porto
Cabedelo
N
Douro
2
1+1B
Intertidal zone
LOEBLICH and TAPPAN (1988). The rippled smearslide technique
was used for calcareous nannoplankton analyses based on the
methodology described by CACHÃO and MOITA (2000). A
cococolith abundace index (CAI) was determined as a count of all
coccoliths present in a 30mm row of the smearlide. Sample
treatment for palynological analysis was based on the
methodology reported by De VERNAL et al. (1996) partly modified
by the “Département de Géologie et Oceanographie (DGO) Université Bordeaux I”. Pollen identification was based on REILLE
(1992) Atla’s as well as on pollen reference collection in the
DGO. At least 350 pollen grains (excluding aquatic plants and
spores) and 100 Lycopodium grains (exotic pollen introduced
during the laboratory proceedings) where counted in each of the
77 samples analysed.
RESULTS
Figure 1: Location of Douro estuary and coring sites.
After visual description the cores were sub-sampled for several
analyses, which included sedimentology (texture, mineralogy of
fine-<63μm – and coarse- >63μm- fractions, organic matter,
carbonate
and
water
content),
geochemistry
and
micropaleontology (foraminifera, calcareous nannoplankton and
pollen associations). Nine organic sediment samples were dated
by 14C AMS at Beta Analytic Inc., USA (Fig.2).
Textural analysis was performed by means of the traditional
sieving methods and gravel-free sediments were classified
according to FLEMMING (2000). The mineralogy of sand was
determined by means of a stereomicroscope observation following
DIAS (1987). Microscope counting of the heavy minerals species,
separated using Na-polytungstate in the medium to very fine
fractions, was made according to the line method (GALEHOUSE,
1969). The study of gravel followed the methodology suggested
by DOBBKINS and FOLK (1970). The organic matter (OM) and
carbonates contents have been determined following CRAFT et al.
(1991) and HULSEMAN (1966), respectively. Mineralogical
analysis of both the fine (<63 µm) and clay (<2 µm) fractions was
carried out using X- Ray Diffraction (XRD) following ROCHA
(1993). The semiquantitative determination of minerals by XRD,
in disaggregated material (<63μm) and in oriented aggregates
(<2μm), followed criteria recommended by BARAHONA (1974),
SCHULTZ (1964), THOREZ (1976), MELLINGER (1979) and PEVEAR
and MUMPTON (1989). For the fraction <63μm, a comparative
analysis of quartz, feldspars and micas content, as well as some
other ratios – FD/CD (fine detritals/coarse detritals) and C/D
(carbonates/detritals) were carried out according to VIDINHA et al.
(1998); for the clay fraction (<2μm), illite, chlorite, kaolinite and
smectite contents and the kaolinite/illite ratio (K/I) were
compared.
Elemental analyses were carried out for major, minor and trace
elements by Energy-Dispersive X-Ray Fluorescence Spectrometry
(EDXRF) in the fraction <63µm of forty samples of core 2.
Sediment portions of about 1.5-2.5g of material were
homogenised and dried at 110°C for 24 hr. The homogenised
material was mixed with an organic binder and pressed into pellets
for analysis. Detailed sampling preparation techniques and
analytical procedures have been published elsewhere by ARAÚJO
et al. (2002).
Foraminiferal analysis was undertaken in the coarse fraction and
a minimum of 100 individuals were counted in each of the 105
studied samples. The ELLIS and MESSINA (1995) catalogue was
used in the species recognition and the taxonomy followed
Sedimentological analysis
According to the sedimentological characteristics, the
sedimentary record of the Douro estuary may be divided in four
main units (Fig.2):
SED1– from 13730 to 10310BP – this unit is present only in
core 2; it consists of slightly muddy sand with some muddy sand
laminae at the bottom layers. It has low contents of fine sediment
(usually less than 25%), OM (0.3-2.7%) and carbonates (00.35%).
SED2 – from 10310 till 5750BP- this unit is present in cores 2
and 1 B, although its topmost section has been preserved only in
core 1B. In core 2, the base of this unit consists of thin
interbedded layers of sandy mud and muddy sand, followed by
slightly sandy mud and muddy sand sections; samples are in
general characterized by high contents of fine fraction, reaching
91% of the total sample at the middle section and decreasing
upwards. OM content, varying between 3 to 11%, mimics the
variation found in the fine fraction. In core 1B, this unit is
represented by alternated layers of sandy mud and muddy sand;
fine fraction contents oscillates between low (1.4-23%) and high
values (47-79%); once more OM matches the oscillations found in
the mud content, varying in general between 2 and 6%, with
extremes of 0.8 to 18%. Some peaks of carbonates can be
observed in the upper part of core 2 and in core 1B, mostly related
with biogenic sand.
SED3 – from post 5750BP and prior to 1580BP - this unit was
found in all cores and is made of gravel and gravelly sand with a
small contribution of mud (0.1-2.5%); OM (0.1-2.9%) and
carbonate (0.2-5.3%) are consistently low.
SED4 – from 1580BP till present - this unit was retrieved in
cores 2 and 1 and it corresponds to a sandy unit capping the
gravel. In core 2, this unit is homogeneous and essentially made
of clean minerogenic sand, while in core 1 there are some
interbedded muddy sand layers, at the bottom and top of the unit;
in these layers, both the fine fraction and OM, acquire significant
importance varying between 40-89% and 5-8%, respectively.
Sand composition
The terrigenous sand component is generally constituted by
mica, quartz and in small amounts by heavy minerals.
The heavy mineral content in the medium to very fine sand is
mainly composed by biotite, amphibole, andalusite, tourmaline,
garnet and apatite among others (FRADIQUE et al., this issue). This
mineral suite reveals a provenance signal related to the igneous
and metamorphic Douro basin bedrock.
The basal levels of SED 1 show a heavy mineral suite resulting
from the direct erosion of the bedrock being mainly composed by
biotite (20%), andalusite (27%) and apatite (6%). The majority of
the mineral grains have euhedral forms, reflecting an incipient
transport. The medium and upper levels of this unit are basically
composed by biotite (40%), amphibole (12%) andalusite (6%),
tourmaline and garnet (both with 3%.).
In SED 2 unit the heavy mineral assemblage consists of biotite
(45%), amphibole (10%), andalusite (4%), tourmaline and garnet
(together with 3%). The levels with high (>50%) content of biotite
match the muddy sand/sandy mud sequences in this unit.
In SED 3, heavy minerals appear only in core 2 (less coarse),
revealing the presence of biotite (49%), amphibole (7%) and
rounded grains of garnet (3%), andalusite (4%) and tourmaline
(2%).
The SED 4 unit has a heavy mineral suite composed by biotite
(32%), amphibole (18%) and andalusite (5%). It should be
referred the occurrence of a distinct heavy mineral suite that
characterizes the upper layers of SED4 in cores 1 and 2. This
distinct suite was used by FRADIQUE et al. (this issue) to define a
sub-unit - SED 4A. It is composed of rounded grains of garnet
(15%), andalusite (13%), tourmaline (8%) and staurolite (5%).
The vertical abundance of marine biogenic components (namely
molluscs, foraminifera, equinoderms, etc.) and plant roots have
opposite behaviour: the former are totally absent in SED1 and in
the lower half of SED 2 (in core 2) while the latter are abundant,
reaching maxima of 34.5% (Fig. 2); further upcore in SED 2, plant
remains are virtually absent and marine biogenic components are
abundant. The molluscs are the most represented biogenic marine
group: they become more important when root content is less
significant (in core 2). Higher values of molluscs usually associate
with maxima in benthic foraminifera and “other biogenics” (e.g.
between -28.11 and -27.45m and at -11.69m). A marked increase
in biogenic component is verified from -8.69m till -8.20m (Fig.2).
SED3 and SED4 are virtually barren in these components: only
a discreet presence of molluscs at -1.69m is verified (in core 1).
Mineralogical
fractions
composition
of
fine
and
clay
Mineralogical characteristics of fine and clay fractions are quite
different in the several sedimentary units.
SED1 unit may be divided in three subzones:
- from the base until -37.66m, the fine fraction mineralogy is
characterized by an almost silicilastic association, constituted by
quartz and feldspars (and therefore, denoting a low mineralogical
maturity), and by a discrete presence of phyllosilicates and rare
carbonates. The clay minerals association is rich in kaolinite and
illite (but showing high K/I ratio), although with a relatively high
expression of smectite or chlorite (Fig.2);
- between -37.66m and -34.07m, quartz increases
simultaneously with feldspars decrease (and a discrete increment
in phyllosilicates) whereas clay fraction is characterized by a
decrease of kaolinite (which leads to the predominance of illite)
and by a much higher increase of chlorite in relation to smectite;
- between -34.07m and -31.16m, a relative increase in quartz
related to the decrease of phylosilicates and feldspars can be
observed. In clay fraction, kaolinite drastically decrease, illite
(very degraded) becomes much more important and smectite
acquire higher importance than chlorite (rare and even absent in
some samples).
This last subzone continues through the lowest layers of SED2
(core2). This unit is mainly characterized (from -31.16m to
-20.66m) by the gradual increase of feldspars and the presence of
micas (and more discreetly, carbonates) in the fine fraction, and by
the decrease of illite (now more ordered) in relation to kaolinite,
whereas smectite decrease comparatively to chlorite. After
-22.74m, C/D ratio increases and K/I decreases. In core 1B, SED2
is characterized by the presence of some detrital minerals such as
quartz, feldspars and phyllossilicates, while the predominant
mineral in the clay fraction is illite, followed by kaolinite
discretely accompanied by chlorite. Smectite gradually increases
in relation to the progressive decrease of chlorite.
SED3 is characterized by a large increase in quartz (particularly
evident for core 2) and a very discrete decrease of chlorite in clay
fraction (core 2); kaolinite increases and smectite decreases,
relatively to SED2 unit, more clearly at the bottom of this unit
(particularly for Core 1) but these tendencies reverse topwards
(core 1) (Fig.2).
SED4 contains the lowest values (almost null) of C/D ratio and
the highest values of FD/CD ratio. In the clay fractions, kaolinite
decreases in relation to illite whereas smectite increases in relation
to chlorite. Illite is once again predominant in a relatively complex
assemblage, illite plus kaolinite plus chlorite plus vermiculite plus
smectite; topwards (core 1), smectite decreases very significantly,
at the same time that other clay minerals increase – such as
vermiculite, mixed-layers and chlorite.
Geochemical analysis
The elemental analysis of core 2 shows that SED1 is
characterized by higher values of Al, K, Si, Ti and Zr, comparable
to the continental crust composition, which can be attributed to
continental input (Fig.2). The systematically low Ca (~0.2-0.5%)
and Sr contents are due to the widespread occurrence of weathered
granitic rocks in the region. However, a significant and
unexpected increase in Ca, Sr (non-common in sediments with a
granitic origin) and Zr was observed at the base of SED1, which is
probably related to some local enrichment in specific minerals (e.g
apatite) in non-weathered basement rock. In particularly, Al, K,
Rb contents should be related to the existence of feldspars and/or
micas of granitic rocks. Besides, the gradually decrease of these
elements till the middle of SED2, negatively correlated with Si
(that increases upwards -32.56m) can be associated with the
decrease in terrigenous input (in agreement with other proxies). It
is worth noting the “anomalous” increasing concentrations in Ca
and Sr found in the upper sections of core 2 (from -25.36m
upwards), probably due to the presence of biogenic marine
organisms.
Micropaleontological analysis
Foraminiferal assemblages mainly record estuarine/brackish,
estuary mouth/marine and outer shelf/slope environments (Fig.2).
The earliest foraminifera occurrence is noticed at -27.77/
-27.66m, with abundant Haynesina germanica, followed by
Elphidium gunteri, Cibicides lobatulus and Amonia spp., which
suggest a brackish environment. After a barren interval of circa
2.5m, foraminifera reappear in the fossil record with H.germanica
(with lower abundance than the levels cited above) or E.gunteri,
followed by A. tepida and some species typical of estuary mouth
and of major depths, suggesting a brackish intertidal, low salinity
estuarine environment. The facies then changes to a subtidal
environment with normal to moderate (-22.39 and -21.48m), or
moderate to low salinity (-21.26 and -20.53m), typical of an
estuarine mouth under marine influence as it is suggested by the
presence of H. germanica E. gunteri and A. tepida occurring
conjointly with Cassidulina laevigata, C. lobalutus and
Gavelinopsis praegeri.
Fine sediments
(<63μm) ( %)
Core 1
0
50
Fine fra ction
mineralogy (%)
100 0
50
100
Clay fraction
mineralogy (%)
0
35
Equinoderms
(%)
1
70 0
Molluscs
( %)
0
Foramini fers ( %)/
Environment type
50
0
50
Nannoplankton
(CAI)
100
0
5500
2.8
1.8
SED4
1110±40BP
0.8
(-0.65m)
-0.2
1580±40BP
-1.2
(-1.80m)
0
300
E
-2.2
-3.2
SED3
D
-4.2
Core 1B
5750±40BP
(-7.15m)
6050±60BP
(-8.34m)
-5.2
-6.2
-7.2
-8.2
0
100
C
-9.2
-10.2
SED2
-11.2
-12.2
B
-13.2
9490±60BP
(-13.91m)
-14.2
-15.2
-16.2
Core 2
0
50
Fine fra ction
Clay fraction
Zr/100ppm
P l. roots
mineralogy (%)
mineralogy (%)
(%)
50
100 0
35
70 0
12 0
30
100 0
Molluscs
(%)
Fine sediments
(<63μm) ( %)
Foramini fers ( %)/ Nanno.
Environt. type
(CAI)
50 0 50 100 0 60 120
0
-15,7
SED4
E
-16,7
-17,7
-18,7
SED3
8930±60BP
(-20.59m)
D
-19,7
-20,7
-21,7
-22,7
9230±50BP
(-22.75m)
-23,7
B
-24,7
-25,7
SED2
-26,7
-27,7
-28,7
-29,7
-30,7
10310±80BP
(-32.33m)
-31,7
-32,7
-33,7
SED1
A
-34,7
-35,7
-36,7
-37,7
13730±90BP
(-39.07m)
-38,7
-39,7
Fine fra ction mineral .
Gravel
Gravely sand
Sand
Slightly muddy sand
Muddy sand
Sandy mud
Slightly sandy mud
Weathering granite
Quartz
Feldspars
Phyllosilicates
Carbonates
Environments Types
Clay mineralogy
Illite
Kaolinite
Chlorite
Smectite
Vermiculite
Interst.
Foramini fera
Estuary/Brackish
Estuary mouth/Marine
Shelf/Slope
A – Continental
B- Marine/Continental
C- Marine
D- Barrier
E- Barrier/Back Barrier
Figure 2– Synthetic results of the analysed cores: lithostratigraphic units (Sed1, SED2, …SED4); fine sediments content, fine and clay
fraction mineralogy, biogenic sand component (equinoderms, molluscs, plant roots); foraminifera assemblage and nannoplankton
“Content Abundance Index” (CAI). A, B,..,E correspond to the defined environmental types.
Between -12.7 and -9.38m, a dominance of H. germanica,
followed by E. gunteri, A.tepida, C.praegeri and C. lobatulus,
indicates a brackish intertidal estuarine environment, of low, and
punctually, moderate salinity; between -9.26 and -8.85m, C.
lobalutus and G.praegeri become dominant, probably in relation
with an intertidal-subtidal estuarine mouth environment with
normal to moderate salinity and clear marine influence (Fig.2).
It is only upwards -8.35m and till the end of SED2 that some
middle continental shelf species as Cassidulina crassarossencis
Globocassidulina subglobosa and Bolivina difformis become more
important. On the other hand, the dominance of C. lobalutus,
G..praegeri and B.pseudoplicata, together with the low abundance
of H. germanica (<5%) and the absence of A. tepida and E.
gunteri indicate maximum marine influence in the estuarine
mouth.
SED3 is barren in foraminifera, and SED4 of core 1 has only
one layer (-1.69m) with foraminifera content.
In general, the estuary mouth/marine and outer shelf
associations increase in SED2 unit (core 2 and 1B). However, this
increase is not uniform, but characterized by some peaks and
interruptions, the former being synchronous with cocoliths peaks,
indicating an increase of marine influence.
In fact, the variations in cocoliths in the SED2 topmost layers
are congruent with the environmental changes deduced from
estuary mouth/marine and outer shelf/slope foraminifera
associations. The earliest occurrence of calcareous nannoplankton
is at -22.74m (which corresponds an age of circa 9320BP). In the
lowermost layers of SED2 in core 1B, a discreet and initially
interrupted presence of coccoliths, dominated by Gephyrocapsa
oceanica and Emiliana huxleyi, suggests an environment with
ephemeral marine influence. Only in the top levels of this unit
(after –8.34m and 6050 BP), the coccolith abundance (>5,000
CAI) and diversity, is compatible with a clear marine environment
in which Helicosphaera carteri, E. huxleyi, Gephyrocapsa
ericsonii, G. oceanica and G. muellerae dominate together with
Coccolithus pelagicus in lesser degree (Fig.2). H. carteri can
exceed 1,000 CAI, equivalent to more than 80% of the
assemblage.
Pollen analysis
Pollen analysis from core 1B, shows the presence of a PinusQuercus-forest with Alnus and some Poaceae-Ericacea open
communities, in the Douro basin, between 9490 BP and 5750 BP.
This forest reflects a temperate and humid climate in the first part
of the Holocene. Pollen preserved in this unit was transported by
the river (NAUGHTON, 2002; NAUGHTON et al., 2002 and
NAUGHTON et al., submitted).
In core 1 the pollen grains are only present in the muddy sand
layers of SED4. These layers are dominated by pollen of
Ericaceae, Poaceae and Asteraceae. This pollen assemblage
represents the signature of the local vegetation characterising
nowadays the surroundings of the coring point.
DISCUSSION – EVOLUTION MODEL
Overall results from this study allowed to distinguish several
environmental changes related with sea-level rise (in particularly,
with the Holocene transgression), regional climate changes
occurred since the Lateglacial and some local forcing factors.
Continental environment (A) – circa 13730-9834
BP
The basal sedimentary record indicates a continental
environment with terrestrial facies that includes SED1 and the
lower half of SED2 (Fig.2). This terrestrial signature is supported
by the total absence of biogenic sand components (namely
molluscs), the significant abundance of plant remains and the total
absence of foraminifera and calcareous nannoplankton. The suite
of heavy minerals, together with the high contents of Al, Ca, Sr
and Zr and the clay mineral association corroborates this
interpretation suggesting evidences of direct erosion of outcroping
basement. This section seems to have been deposited in a
changing climate: it evolved from temperate to subtropical
conditions (from base to -37.66m) to milder ones (-37.66m/
-34.07m) and finally to temperate conditions (but with colder
periods) with an increase of precipitation (-34.07/-31.16m).
SED1 sequence probably corresponds to a typical fluvial
environment, in the dependence of the Douro river. It was
deposited when the shoreline was far away from the coring sites:
circa 13km (between 13-11ka; sea level at -30m) and 20km
(between 11-10ka; sea level at -60m) according to DIAS (1987).
The sedimentation becomes finer in the lower part of SED2 unit,
probably in response to the approximation of the base level.
Marine/Continental
6050BP
environment
(B):
9834-
The following facies is characterized by alternating marine/
continental influences and corresponds to the upper half of SED 2
(in core 2) and SED2 (in core 1B) (Fig.2). This facies is
characterized by increasing marine influence, as suggested by the
general increase of estuary mouth/marine and continental shelf
foraminifera (especially expressed in core 1B). However, drastic
decrease in abundance or even interruption in the foraminiferal
record, molluscs and nannoplankton can be observed at specific
levels (Fig.2). These perturbations, are simultaneous with
augmentation in terrigenous indicators, especially plant remains.
Thus, these intervals were attributed to continental inputs, which
seem to have been short-lived (some of them may correspond to
intervals as short as 130-200 years, considering the radiocarbon
dates and a uniform sedimentation rate) and disturbed the
transgressive trend. They can represent either regressive periods
or, more probably, episodes of important fluvial sediment input.
The earliest marine signal is represented by a single ephemeral
episode, with brackish foraminifera assemblage and some peaks of
molluscs and “other biogenics”, noticed between -28.11m and
-27.45m and corresponding to an interpolated time period between
circa 9834 and 9756 BP. However, it is after a threshold circa
-25m that the signal of sea level incursions persists, as shown by
the abundance and associations of molluscs and foraminifera.
The marine signature is also supported by the increase in C/D
ratio and decrease in K/I, particularly after -22.74m, when
augmentations in Ca and Sr are also observed.
Besides, the presence of nannoplankton at this specific level
(corresponding to circa 9230 BP) indicate that sea level had
already exceeded this elevation by this time. According to DIAS
(1987), after 11-10ka, when sea level was circa -60m, a rapid rise
took place, reaching approximately -20m at 8ka. However, this
new data seem to indicate that this level was reached at an earlier
stage (circa 9ka), which may imply a higher sea level rise rate than
previously proposed.
The pollen assemblages are in agreement with the climatic
amelioration that characterizes the Holocene as it was deduced by
pollen analysis. In particular, XRD data of core 2 put in evidence a
climate evolution towards more temperate conditions, with less
precipitation and, therefore, less favorable to hydrolysis
development in source areas. In core 1B, hydrolysis seems to
increase progressively, towards a warmer and more humid
climate, with seasonal contrast, simultaneously with less
hydrodinamism.
Marine environment (C): – 6050-5750BP
A clear marine environment upwards -8.35m is very well
expressed (and corresponds to the topmost layers of core 1B),
through the foraminifera assemblage (effective presence of
continental shelf species and minor importance of brackish
environment species), nannoplankton diversity and abundance
(Fig.2) and by the evident increase of C/D ratio. These data
confirms that the coastline stood landward of the core location,
when sea level was circa 7m below present day level, representing
thus the maximum of the Holocene transgression. However, core
1B site was probably related to a partially confined environment,
probably by the precocious development of a sandy bar as it seems
to be suggested by the observed abnormal opportunistic
development of H.carteri which denote some ecological
particularities (CACHÃO et al., 2002).
This environment can be considered an extension of the
preceding one, and some of its characteristics, as the climatic
ones, remained unchanged from the B to the C environment.
Barrier environment (D): post 5750 – before 1580
BP
After 5750 BP, a clear change in the environment is expressed
by the deposition of a gravel layer - SED3 (Fig.2). This unit seems
to have had a fluvial origin and subsequently a littoral reworking
(NAUGHTON, 2002; NAUGHTON et al., submitted), which seems to
be supported by the rounded, spherical and ellipsoidal grains of
garnet, andalusite and tourmaline in this unit of Core 2. This
gravel layer may represent a gravel barrier, which formed post
5750 and before 1580 BP (NAUGHTON, 2002, NAUGHTON et al.,
submitted) and thus, contemporaneous of similar features (though
sandy) found in the SW Portuguese coastal lagoons of Albufeira,
Melides and Santo André (FREITAS and ANDRADE, 2001).
The extension of this gravel barrier allowed the northward
relocation of the Douro channel (NAUGHTON, 2002, NAUGHTON et
al., submitted), leaving its former position fossilized as a
paleovalley located south of the present-day talweg (CARVALHO
and ROSA, 1988). This could be responsible for the observed
change in the pollen signature from regional to local before 5750
and post 1580BP, respectively.
Climatic proxies seem to indicate an increase in temperature
and precipitation, favouring a strong hydrolysis in source areas
and an intense fluvial transport as it was deduced from
mineralogical data.
Barrier/back-barrier environment (E): 1580BP –
present day
A barrier/back-barrier environment is evident in SED 4, with a
brackish signal suggested by micropaleontological data and a
small percentage (<1% molluscs) found in muddy layers at its
bottom (-1.69m) (Fig.2). With the exception of this episode, no
biogenic sand components have been detected further upcore. This
facies includes a significative percentage of amphibole in the
heavy mineral suite of the top of the unit, which is probably
related with the amphibolitic outcrops in the southern area of
Douro basin.
The accumulation of the estuarine spit and sedimentation in the
estuarine domain sheltered by the barrier, progressed at a high
rate, allowing the marine signal to fade away in the topmost
section of the sedimentary record of this place. Therefore, a forced
regressive facies was constructed here in spite of a persistent
eustatic positive trend.
XRD data seems to show the existence of hydrolysis conditions
with seasonal contrast and a loss of fluvial hydrodinamism.
The conceptual model out coming from these results follows in
the essential the one accepted for the southwestern Portuguese
lagoons of Santo André and Melides (FREITAS et al., 2003). Since
the Lateglacial until the Early Holocene the sedimentary record of
these lowlands shows a pronounced terrestrial character (FREITAS
et al., 2003). Sediments are generally coarse-grained
characterizing detrital fluvial input, and bearing a strong
provenance signal from the watershed. In contrast, the overlying
deposits show a dominant marine influence, the sediment being in
general finer, enriched in OM bioclastic particles (FREITAS op.
cit.). This influence reached a maximum circa 5000-5500 BP,
when the deceleration of sea-level rise rate allowed the
emplacement of detrital barriers and the establishment of restricted
environments, which evolved as such until present.
CONCLUSIONS
The sediments of the Douro estuary were deposited through a
succession of different environments, mostly related with changes
in sea-level rise, specially, with the Holocene transgression,
regional climate and local forcing factors that occurred since the
Late glacial. The proposed evolution model agrees with the
existing one for the southwestern Portuguese lagoons of Santo
André and Melides (FREITAS et al., 2003), reinforcing some
original aspects that seem to have a regional meaning.
In fact, since 13730 and till 10310 BP, the sedimentary record
shows a terrestrial character followed by a marine environment,
characterized by an alternating marine/continental facies (103106050BP). However, a clear marine influence expressing the
maximum transgression event is possible to observe only after
6050 and till 5750BP. After 5750 BP, the establishment of a
gravel barrier followed by a sandy barrier development allowed
partial sheltering of the estuary. Thus, the post-Late glacial
evolution of the Douro estuary includes a transgressive facies
followed by a regressive one. It represents therefore a complete
sedimentary cycle in spite of a positive eustatic signal that
persisted throughout almost this entire period.
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ACKNOWLEDGEMENTS
This work was undertaken as a part of the project ENVICHANGES-PLE/12/00 funded by Fundação para a Ciência e a
Tecnologia (Portugal). The authors thank to C. Andrade for his
field work assistance and enlightening discussions and R. Taborda
and M. Fernanda Sanchez Gõni for the carefully revision of this
paper.

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