Printed in the Netherlands. Organo-geochemical and stable isotope indicators of environmental change in a marl lake, Malham Tarn, North Yorkshire, U.K.. 1,. 2.
Journal of Paleolimnology 28: 403–417, 2002. 2002 Kluwer Academic Publishers. Printed in the Netherlands.
403
Organo-geochemical and stable isotope indicators of environmental change in a marl lake, Malham Tarn, North Yorkshire, U.K. ˜ 1, *, B. Spiro 2 , A. Pentecost 3 , A. Kim 4 and P. Coletta 3 R. Nunez 1
´ ´ ´ Experimental del Zaidın ´ ( C.S.I.C.), Departamento de Biogeoquımica de Isotopos Estables, Estacion Profesor Albareda 1, Granada 18008, Spain; 2 Department of Mineralogy, The Natural History Museum, Cromwell Road, London SW7 5 BD, United Kingdom; 3 Division of Life Sciences, Franklin-Wilkins Building, King’ s College London, 150 Stamford Street, London SE19 NN, United Kingdom; 4 British Geological Survey, Analytical and Regional Geochemistry Unit, Nicker Hill, Keyworth, Nottingham NG12 5 GG, United Kingdom; * Author for correspondence (e-mail: rafael.nugnez@ eez.csic.es) Received 14 January 2002; accepted in revised form 11 June 2002
Key words: Biomarkers, 13 C, Chara, Last glacial Holocene, Marl lake, Molecular d 13 C
Abstract Sediments of the marl lake Malham Tarn located in NW England preserve an environmental record since 12 Ka. Eight Holocene pollen zones were identified, and the d 13 C of total organic carbon (TOC) shows three stratigraphic divisions. The basal clay unit and overlaying sand / clay / marl unit have d 13 C of 224‰ which decreases at the base of the principal marl unit to a mean value around 230‰, whilst the topmost black marl unit d 13 C increases to 228‰ at the surface. Representative samples of these units were selected for analysis of n-alkanes and n-fatty acids and their d 13 C. Samples of modern Chara and peat were analysed for comparison. The clay unit has a minor contribution of redeposited mature organic matter and autochthonous algae, the marl unit a high contribution of Chara, and the dark marl unit has a high contribution from higher plants. Compound-specific d 13 C reveals systematic differences between alkanes and fatty acids of different chain length. The major shift in d 13 C in the short and medium chain fatty acids are probably due to the decreasing influence of carbonate rock flour as source of DIC. The major shift in d 13 C in the long chain n-fatty acids and n-alkanes could reflect the lower atmospheric CO 2 concentration at Last Glacial. The negative shift of short chain fatty acids in organic rich dark marls reflects introduction of detrital peat into the lake. The d 13 C results show a dramatic change from dominance of autochthonous plus eroded sources up to Pollen Zone IV, then slow colonisation of the hinterland by higher plants, followed by constant Chara contributions throughout the deposition of the marl, and a further increase of higher plant material after the rise in water level in 1791.
Introduction Lacustrine organic matter has proved to be a powerful archive of environmental and climate records. Organic matter preserved in lake sediments is a mixture of the remains of terrestrial and aquatic organisms (Meyers and Ishiwatari 1993a, 1993b). Bulk carbon isotope composition of lacustrine organic remain, reflects variation in limnological processes and has been used for paleoclimatic and paleoenvironmental purposes in many lakes worldwide (Hakansson 1985; Hammarlund 1992; Street-Perrott et al. 1998; Mayer
and Schwark 1999). But it represents mixed isotope signals from land plants, aquatic plants and microbes. It is therefore difficult to distinguish isotopic variations that reflect changes in the biogenic constituents from those that may be truly indicative of an environmental influence. Using the molecular carbonisotope composition measured by gas chromatography isotope ratio mass spectrometry (GC-IRMS) of biomarkers, it is possible to refine the interpretation of the bulk carbon isotope composition (Freeman et al. 1990; Rieley et al. 1991; Bird et al. 1995; Huang et al. 1999; Brincat et al. 2000). Different lipid biomarkers
404 represent different types of organic matter and can be used to fingerprint their different sources. Long chain, normally .C 25 , n-alkyl lipids such as n-alkanes, nalkanols and n-fatty acids, occur in leaf waxes of terrestrial plants (e.g., Eglinton and Calvin (1967), Eglinton and Hamilton (1967), Kolattukudy et al. (1976)). Shorter n-alkyl lipids are abundant in aquatic plants (Barnes and Barnes 1978; Viso et al. 1993; Ficken et al. 2000). Carbon isotope composition of lipid biomarkers can then be used for tracing carbon sources in ancient sediments (Freeman et al. 1990; Collister et al. 1994; Ficken et al. 2000), and for palaeoenvironmental interpretations (Street-Perrott et al. 1997; Huang et al. 1999; Brincat et al. 2000). The Last Glacial–Holocene transition shows a major carbon isotope shift in TOC in lake sediments worldwide, with values up to 20‰ lower in the Holocene than in the Last Glacial. By using the compoundspecific carbon-isotope analysis, these differences have been attributed to the change in the relative distribution of C4 plants in high altitude tropical lake sediments. The C4 plants would have been selectively advantaged compared with C3 plants by drier climates and lower CO 2 concentrations (Street-Perrott et al. 1997; Huang et al. 1999). However very few studies have been carried out on high latitude lakes. For example in lake Baikal, Brincat et al. (2000) found no changes in the carbon isotope composition of long chain n-alkanes throughout this period, while the same samples displayed significant changes in bulk carbon isotope composition. The authors show that this stability of n-alkanes d 13 C composition, with values typical of C3 plants, is unlikely to be attributed to an expansion of C4 type vegetation in the Baikal watershed as had been previously published. It is therefore interesting to have more data about high latitude lakes, which display strong shift of TOC d 13 C in the Last Glacial–Holocene transition but with scarce possibilities of having C4 vegetation. Lake sediments from different locations in UK such as Gransmoon and Llanilid (Turney et al. 1997) and Lundin Tower (Whittington et al. 1996) show variation in the Last Glacial–Holocene transition, with d 13 C higher by 3 to 10‰ in the Last Glacial than in Holocene. However, to the best of our knowledge, no compound-specific isotope analysis of lacustrine organic matter has been performed in any of them. Malham Tarn, a small shallow lake in Yorkshire, U.K., is one of the few upland marl lakes in the British Isles (Fryer 1991). It possesses phytobenthos dominated by calcite-depositing macrophyte Chara
globularis ssp. Virgata. It is the only lake in Britain for which it is reported that Chara marl accumulated for the past 10,000 years and continues (Pigott and Pigott 1963; Holmes 1965; Coletta et al. 2001). Chara remains are an important source of organic matter in the sediments of marl lakes and were used for palaeoenvironmental interpretations (Jones et al. 1996; Hammarlund et al. 1999). In the present work we have determined the quantitative distribution and Chara carbonate, the distribution of TOC and its d 13 C, and the quantitative distribution of some biomarkers (n-alkanes and n-fatty acids) and their d 13 C from the lake sediments spanning the period Lateglacial to present day. The results have provided information about the palaeoenvironmental changes that occurred in the lake and support the idea that lower CO 2 concentration in last Glacial could have played a major role in the d 13 C of higher plants in high latitudes.
Experimental Malham Tarn (Figure 1) is a small lake situated 8 km east of Settle, North Yorkshire, latitude 548 059 N, longitude 28 109 W at 375m above sea-level on a broad shelf of Great Scar Limestone (Visean) which forms the south-west margin of the Craven Pennine upland. The tarn occupies the deepest part of an irregular depression in the hummocky surface of a layer of glacial drift, which partly fills a more extensive basin situated immediately to the north and on the upthrow side of the North Craven Fault. The original basin corresponds roughly with an inlier of Silurian slates, but is almost completely surrounded by limestone except for the eastward extension of drift over Mastiles and an outlier of grits on Black Hill to the west. The tarn water is derived from a series of springs, which probably originate from the contact between the Silurian slates and the limestone, and most issue at the base of limestone cliffs on the north side of the basin. Two Livingstone piston cores were taken. The first, coded 96 / 2 was taken near mid-lake (Figure 1) and penetrated 6.5m of marl (Nat. grid ref. 34 / 893667). The second, coded 99 / 1 was taken at the western peat margin of the Tarn just onshore (Tarn Moss) (Nat. grid. ref. 34 / 889670). This core penetrated 2m raised peat before entering lake sediments which consisted of about 0.7m marl followed by 1.9m clay. Cores were extruded into polyethylene tubes and kept in a
405
Figure 1. Location map of Malham tarn
cold room at 58C. Edge-cleaned samples of ca 30g wet weight were removed with a steel knife, placed on aluminium discs and freeze-dried at 2208C. Eight samples were selected from the core 96 / 2 and seven from 99 / 1 for biomarker studies. Additionally, one sample of modern Chara globularis was selected for comparison. Several additional samples throughout the two cores were selected for analysis of bulk organic carbon isotope and Chara carbonate content. For biomarker studies, between 0.5g and 20g of freeze-dried sediment were powdered, then extracted ultrasonically using a mix of CH 3 OH:CH 2 Cl 2 (1:1) to obtain free lipids. After centrifugation and concentration, sulphur was removed by adding copper wires, and the solvent dried by passing over anhydrous Na 2 SO 4 . Total extracts were split into an acid and neutral fraction by solid phase extraction with Aminopropyl Bond Elute. For each sample, a new column was prewashed with CH 2 Cl 2 and 2:1 CH 2 Cl 2 :isopropanol. Each column quantitatively retains acids when total extracts are flushed through with 2:1 CH 2 Cl 2 :isopropanol. The acid fraction was subsequently recovered with 2% acetic acid in ether and then methylated with 10% BF 3 -CH 3 OH overnight at room temperature. The neutral fraction was fractionated into aliphatic hydrocarbons, aromatic hydrocarbons and polar compounds by thin layer chromatography (silica gel 60, 0.25mm thick: solvent n-hexane). Identification and quantification of the abundance of individual n-alkanes within the aliphatic fraction was accomplished using a GC and GC-MS with internal standards. Gas chromatography (GC)
analyses were carried out on a Fisons GC8000 GC fitted with split / splitless injector and FID. A Cpsil5CB (Chrompack) fused silica capillary column (60m 3 0.32mm; 0.12mm film thickness) was used. Oven temperature was held at 408C for 1min, ramped at 108C min 21 to 3208C and held there for 25min, with He as carrier gas. Squalene was used as internal standard. Gas chromatography-mass spectrometry (GC-MS) analyses were performed on a Fisons GC8000 (oncolumn injection, 70eV EI) interfaced directly with a Fisons MD800 Quadrupole mass spectrometer. The column, temperature program and carrier gas were the same as for the GC analyses. Gas chromatography-combustion-isotope ratio mass spectrometry (GC-IRMS) analyses were performed using a GC8000 (split / splitless injector) attached to an OPTIMA isotope ratio mass spectrometer via a combustion interface (8508C) consisting of a 60cm quartz tube (0.65mm ID) filled with copper oxide. A liquid nitrogen cold-trap was used for removing water. The column, carrier gas and temperature program were the same as for the GC analyses. Each sample was run in triplicate to ensure reproducibility (normally better than 60.5‰, 1s). All carbon isotopic ratios are expressed as ‰ relative to the VPDB standard. Samples were concentrated so that a peak height of at least 1 3 10 212 A. was achieved for the compounds of interest. Several pulses of CO 2 with a precalibrated isotopic composition were introduced into the mass spectrometer before and after elution of the n-alkanes or n-fatty
406 acids. Squalene previously analyzed (off-line combustion method) for 13 C isotope composition was coinjected as internal standard. Between each two injections of a sample, a standard mix of C 20 , C 25 and C 30 n-alkanes with known carbon isotope composition (Chiron AS), were measured to test the system stability. The C 22 n-fatty acid methyl ester of several samples was not measured due to a plasticiser (GC-MS m / z 149) contaminant that coeluted with it. Prior to the analysis, n-alkanes / alkenes were isolated from the aliphatic fraction by urea adduction as described by Ficken et al. (1998). Due to the derivatization process a calibration of d 13 C values for fatty acids was needed. This was achieved by using a mass balance relationship based on the off line determination of d 13 C of the BF 3 -CH 3 OH, as reported by Abrajano et al. (1994), Duan et al. (1997). Bulk d 13 C of total organic carbon (TOC) of the same samples used for the lipid extraction and others from the same cores, were measured by automated on-line combustion (Carlo Erba NA1500) followed by dual inlet isotope ratio mass spectrometry (OPTIMA). Carbonates were removed by prior washing with diluted (10%v / v) HCl. The results are expressed as ‰ relative to the VPDB standard. Reproducibility for duplicate analysis was 60.1‰ (1s). To determine the quantity of organic matter, ca. 0.5g of sediment dried at 1058C was weighed accurately, the carbonate dissolved in diluted (10%v / v) HCl, at 208C and ca. 3h., filtered through weighed and washed glass microfibre filters (Whatman GF / C), and after drying, the residue was analysed by thermal analysis at (5508C). Organic matter was obtained by difference. To determine the abundance of Chara carbonates, samples of ca 20g dry wt were sieved at 500mm and the resulting fractions dried and weighed. The coarse fraction was resuspended in water, placed in a petridish fitted with grid lines and examined under a binocular microscope (310). In each sample the relative proportion of the Chara incrustations (carbonates) was estimated and converted to approximate percentage dry weight of the sample.
Results Sedimentological description of cores and pollen zone correlation a) Core 96 /2 This core was taken entirely in the marl (Figure 2).
The upper 300mm consists of a brown-black sediment containing about 60% CaCO 3 , 30% organic matter and 20% silt, sand and clay. The dark colouration is enhanced by the presence of eroded peat from the adjacent peat bog and soot particles deposited from the atmosphere over the past 200 years. This layer is believed to begin with the construction of a dam at the lake’s outflow in 1791 (Pentecost 2000) and is termed here the ‘post dam-event marl’. The dam raised the lake water-level by about 1.2m leading to flooding of the lake edge and erosion of Tarn Moss, the large raised bog at the eastern edge of the Tarn. This event partly explains the raised level of silt and clay in the catchment, although this may also have been due to changes in agricultural practice in the 19 th Century. Detailed investigations have shown that about 8,000kg of peat have been eroded from the Moss every year since this event, with much of the material ending up in the Tarn sediments (Pentecost 2000). The contribution of soot particules from air pollution events is not known but is probably low. Below 30cm the marl changes rapidly to a soft, grey-cream deposit (5Y 6 / 2) with occasional Mollusca, Ostracoda and abundant Chara remains. In places a crude lamination is apparent but overall it is a homogeneous deposit containing on average 95% CaCO 3 , with about 3% organic matter and 2% silt. The organic matter fraction contains both aquatic and terrestrial plant material but most of it is highly degraded and difficult to identify. However, remains of Sphagnum are rare, suggesting that erosion from the Moss prior to 1791 was insignificant. A careful search was made of organic Chara remains but it appeared that apart from oogonia, which were abundant at most levels, vegetative remains of Chara were too degraded to allow identification. b) Core 99 /1 This core was taken from Tarn Moss (Figure 1) where 2m raised peat was encountered above the marl. In this core the layer of marl was much thinner, amounting to only 70cm (Figure 2). This marl, which forms a sharp contact with the peat above was pale cream (5Y 7 / 2) grading below into cream-grey (5Y 6 / 2). It was indistinctly laminated and contained similar fossils with a similar gross composition to the pale marl of 96 / 2. Below the marl (70–100cm below the peat), sediments grade into a grey coarse sandy fraction consisting of angular quartz sand, calcareous clay and fragments of limestone and a little charcoal. At 100cm below the peat, this coarse fraction gave way to a stiff grey clay (5Y 4 / 1) that contained about 50% fine-
Figure 2. Summary of the stratigraphical sections of cores 99 / 1 and 96 / 2 and pollen data, Malham Tarn, depth profiles of Total Organic Carbon (TOC), content of carbonate derived from Chara (in wt%), and d 13 C of TOC.
407
408 grained calcium carbonate and 1–2% organic matter. This clay continued to the base of the core. In places it contained grit and small stones (175cm below peat) and in others it was varved (180–190cm). Near the base of the core it became more brown in colour (5Y 4 / 2). Pollen dating We have pollen spectra for 96 / 2 and 99 / 1 (Figure 2). The top 25cm of the in-lake core, 96 / 2, contains eroded peat as indicated above. Pollen analysis suggests that the base of the core penetrates mid-lake Boreal sediment ca. 8 ka in age. Continuous sedimentation is indicated, at about 0.8mm / yr. The shorter core 99 / 1 penetrates 2m ombrotrophic peat overlying 70cm marl with grit and clay below (Figure 2). The peat formed ca 6–7.5 ka according to our pollen analysis and that of Pigott and Pigott (1959). The marl in this core is older than that of 96 / 2 and according to the palynology appears to have been formed during the Upper Dryas cold state and before (Pigott and Pigott 1959; Pentecost et al. 2000). Selected pollen spectra (Alnus and Pinus) are shown for both cores in Figure 2. Alnus pollen is frequent throughout 96 / 2 excepting the base where it becomes scarce, corresponding to pollen zone V or below. In 99 / 1 Alnus pollen is rare or absent throughout. Therefore the marl in 99 / 1 must be at least as old as the base of 96 / 2. Pinus pollen occurs throughout 96 / 2 but is only prominent at the core base. In 99 / 1 Pinus pollen is frequent throughout, becoming dominant in the lower clays. Therefore the base of 96 / 2 must again be at least as old as the 99 / 1 marl but not as old as the clay below. Our total pollen spectra have allowed us to assign Godwin pollen zones (Godwin 1975) to both of these cores except the upper section of 96 / 1 where it was not possible to establish positions of zones VIIb & VIII. This was due to the scarcity of key pollen indicators among the herbs. Thus it appears that the two cores, 96 / 1 and 99 / 1 span most of the postglacial (0–12 ka BP) and the palynology indicates some overlap between them (Figure 2). Distribution of TOC, n-alkanes, n-fatty acids and their d 13 C The distribution and relative abundance of the homologous series of n-fatty acids and n-alkanes are shown in Figure 3 and in Table 1. Those of Chara and peat are given for comparison. The n-alkane distribu-
tion ranges from C 21 (some at C 20 ) to C 33 for most of the samples except for 325(99 / 1) and 425(99 / 1) which have a wider range from C 14 to C 33 . The most abundant homologue changes from one sample to another in the range C 23 to C 29 , but in all samples the range C 25 –C 31 have an odd-over-even predominance with high Carbon Preference Index (CPI) values (Table 1). The CPI ranges from 2.1 up to 62.9. This indicates for most of the samples have a major component of epicuticular leaf wax origin (Eglinton and Hamilton 1967; Collister et al. 1994). Average Chain Length (ACL) values are consistent between 25 and 29 (Table 1). Samples 325(99 / 1) and 425(99 / 1) show a strong peak of pristane (see Figure 5). The 360(99 / 1) sample maximizes anomalously at C 24 . The n-fatty acid distributions are similar in most of the samples. Normally they show a bimodal distribution and maxima at C 16 or C 18 for the shortest homologues and at C 24 or C 26 for the longest, except Chara which maximises at C 22 . This behaviour has been described as typical of sediments with mixed terrestrial and algal sources of alkyl lipids (Giger et al. 1980; Meyers and Eadie 1993; Ostrom et al. 1998). Some samples contain minor amounts of C 18:1 pointing to an algal origin (Cranwell 1974; Sargent and Whittle 1981; Shameel 1990; Oldenburg et al. 2000). Neither branched nor hydroxy acids were found. The amount of n-alkyl lipids (as mg / gdw and mg / gTOC, see Table 1) decreases with depth from the post damevent marl to the basal marl but increase again strongly in the marl-clay transition. Most of the samples have equal or higher amounts of fatty acids than the n-alkanes. This is in agreement with previous observations made on lacustrine sediments (Ficken et al. 1998, 2000; Huang et al. 1999). TOC concentration remain around 25% in the first 30cm of core 96 / 2. Then decrease to around 5% throughout most of 96 / 2 and 99 / 1, with some larger fluctuations (5–10%) in the lower half of 96 / 2 and in the marl-clay transition of 99 / 1. Downcore the samples show a pattern typical of mixed sources of alkyl lipids. The two end-members would be represented by a) terrestrial long chain n-alkanes and fatty acids such as those of the peat or post dam-event marl, and b) algal short chain and medium chain n-alkanes and fatty acids such as those of Chara. The d 13 C results are shown in Table 2 and Figure 2,4. The d 13 C values of the total organic fraction from the core show three intervals: (1) a slight decrease from 228‰ to 230‰ in the post dam-event marl, (2) constant values throughout the marl, and (3) a signifi-
409
Figure 3. The distribution of n-alkanes and n-fatty acids in the samples from the 99 / 1 and 96 / 2 core. Samples of living Chara and two peat samples from Tarn Moss are reported for reference.
28.4 19.9 16.9 13.9 2.4 4.4 7.3 5.4 18.7 5.1 7.4 4.1 0.8
n.d. n.d. n.d.
0(96/2) 18(96/2) 25(96/2) 30(96/2) 38(96/2) 365(96/2) 500(96/2) 590(96/2) 235(99/1) 260(99/1) 325(99/1) 360(99/1) 425(99/1)
Chara Peat 190(99/1) Peat 180(99/1)
0.9 0.01 0.8
126.4 119.2 15.3 76.2 5.1 3.5 256.7 65.4 99.2 n.d. 15.7 14.6 1.9
161.4 169.9 65.1 55.4 7.2 23.8 459.8 163.6 533.9 n.d. 31.9 4.2 1.4 mg/gdw 0.8 29.0 7.9 0.4 17.6 2.4
235.6 304.6 73.3 58.1 9.7 25.8 319.6 125.3 208.6 n.d. 34.1 4.3 1.4
nC29
0.6 7.3 0.9
355.0 818.9 31.8 49.9 8.0 12.6 8.1 22.9 38.8 n.d. 19.1 4.9 1.1
nC31
1452.7 2258.2 271.1 439.2 48.2 89.5 1674.5 539.5 1514.0 n.d. 163.2 78.3 14.3
Sum mg/gTOC
26.1 27.3 27.2
27.8 28.8 27.6 26.5 27.3 27.9 26.2 26.6 26.5 n.d. 27.3 25.2 26.4
ACL
3.3 18.2 48.7
7.3 7.3 8.6 4.4 9.3 7.4 62.9 39.1 27.4 n.d. 7.0 2.1 4.3
CPI
6.5 87.2 14.0
0.5 1.1 0.2 0.3 0.2 0.2 2.3 1.0 0.8 n.d. 0.2 0.2 0.2
Sum mg/gdw
922.0 82.7 0.0
2472.0 532.0 29.4 59.2 26.4 32.3 832.9 201.3 52.3 6.1 14.4 1.1 0.2 51.0 59.4 0.0
1219.0 593.0 20.7 25.1 9.7 31.3 160.5 29.5 23.3 2.5 6.5 1.6 0.2
nC24
880.0 1882.8 2148.4 1820.3 17.8 17.4 22.1 27.1 5.5 4.3 16.7 42.2 233.9 3075.0 93.4 3056.0 37.5 268.9 6.5 502.0 27.7 1860.0 0.9 7.7 0.02 0.5 mg/gdw 360.0 212.0 94.5 3457.0 0.0 550.9
n-Fatty Acids (mg/gTOC) nC16 nC18 nC22
where i 5 carbon number ranges, Ci5 the concentration of the homologues containing i carbon atoms. CPI: 2 3 odd C25toC33 / (even C24toC32 1 even C26toC34).
(Sum[Ci]3i) ACL (Average Chain Length): ]]] Sum[Ci]
0.8 19.6 1.1
148.7 127.4 34.7 64.7 9.4 9.6 330.0 110.8 388.7 n.d. 20.2 7.2 1.7
n-Alkanes (mg/gTOC) nC23 nC25 nC27
TOC: total organic carbon.
TOC (%)
Sample
0.0 6437.6 2937.9
587.9 1853.4 38.7 30.9 4.9 63.7 2826.0 3652.0 363.9 3963.0 5818.0 36.1 2.0
nC26
Table 1. Summary of quantitative data on n-alkanes and fatty acids from individual samples from the 99 / 1 and 96 / 2 cores from Malham Tarn.
0 3807.6 1482.1
66.5 923.1 20.3 10.3 1.5 26.2 369.0 1815.0 104.5 515.0 629.0 31.6 1.3
nC28
7952.9 8617.5 164.0 207.8 54.3 239.1 7999.8 9217.8 934.0 5056.1 8416.3 87.6 5.2
Sum mg/gTOC
22.5 26.3 26.5
23.9 24.6 25.6 24.9 24.5 25.5 25.1 25.7 25.6 26.0 25.7 26.8 26.6
ACL
1708.4 16553.4 5187.7
19.0 4.3 0.1 0.1 0.2 0.5 10.7 16.5 0.5 9.9 11.4 0.2 0.1
Sum mg/gdw
410
411
Figure 4. Schematic time profiles of d 13 C of TOC and weighted-average d 13 C of n-alkanes, n-fatty acids in the 99 / 1 and 96 / 2 cores from Malham Tarn. Medium chain n-alkanes (C23,C25), long chain n-alkanes (C27, C29, C31), short chain n-fatty acids (C16, C18), medium chain n-fatty acids (C24) and long chain n-fatty acids (C26, C28).
cant increase through the clays to reach values around 222‰ (Figure 2). All these values are typical of C3 plant species that generally exhibit values from 232‰ to 220‰ (Deines 1980; O’Leary 1981). Samples of Chara and peat, chosen for comparison, show values typical of C3 plants too (Table 2), and
agree with those reported previously for peat (Ficken et al. 1998). The d 13 C of n-alkanes range from 224.2‰ to 233.9‰ and the n-fatty acids from 222.1‰ to 237.7‰. With the exception of some n-alkanes and n-fatty acids of clay samples, all of these values point to a contribution of C3 plant tissue,
412 which has typically d 13 C values more negative than 228‰ (Collister et al. 1994; Lockheart et al. 1997). As a general rule, each sample shows more negative d 13 C values (up to 7‰) in the longest chains. This feature has been noted for long chain n-alkanes in lake sediments from Baikal (Siberia, Russia), Nkunga (Mt. Kenya, East Africa) and Priest Pot (U.K.) among others (Rieley et al. 1991; Spooner et al. 1994; Ficken et al. 1998; Brincat et al. 2000) respectively.
organic matter with different proportions of each one through the core. Samples from the post dam-event marl are more enriched in the middle and particularly long chain homologues (C 31 n-alkane and C 24 and C 26 n-fatty acids) indicating a source predominantly of higher plants with a minor contribution from aquatic macrophytes and algae (as indicated by the presence of C 18:1 in sample 0(96 / 2)). This agrees with the fact that an increase in the erosion of peat was observed in the lake since the building of the sluice gate in 1791 (Pentecost 2000) and would explain the high level (25%) of TOC in these samples (see Figure 2). Agreement between the d 13 C values of C 16 , C 18 , C 24 , C 26 and C 28 n-fatty acids in the peat samples and in the post dam-event marl (Table 2) supports this conclusion. In the marl, n-alkane distributions are marked by a higher abundance of mid chain C 23 and C 25 homologues, particularly in the samples 30(96 / 2) and 38(96 / 2), which also show an important contributions of C 16 and C 18 n-fatty acids (Figure 3). Chara remains (Figure 2) are found throughout the 96 / 2 core, but are more abundant around 400mm depth where Chara carbonates make up .30% of the coarse fraction, and other smaller maxima of Chara carbonate occur at 80, 180 and 290cm. These observations suggest that Chara is an important source of organic carbon in the marl samples. The d 13 C values of C 16 and C 18 n-fatty acids in marl samples (Table 2), with the exception of samples 235(99 / 1), 260(99 / 1) (which do not have Chara carbonates), and C 18 of 38(96 / 2), are the most
Discussion Organic matter distribution Organic matter found in lake sediments is a mixture of autochthonous material, algae, aquatic macrophytes and bacteria and allochthonous material, residues of biota from the land surrounding the lake (e.g., Meyers and Ishiwatari (1993a)). Long chain n-alkyl lipids, .C 25 , with high CPI values, have been reported as indicators of input from terrestrial higher plants as they form part of their epicuticular leaf waxes (Eglinton and Hamilton 1967; Kolattukudy et al. 1976). The mid range C 20 to C 24 n-alkyl lipids are more indicative of submerged macrophytes (Cranwell 1984; Ficken et al. 2000) and short ,C 20 chain n-alkyl lipids (saturated and particularly unsaturated) are more indicative of algae (Brooks et al. 1976; Barnes and Barnes 1978; Huang et al. 1999; Oldenburg et al. 2000). The n-alkanes and n-fatty acids distribution in our samples indicates a mixture of different sources of
Table 2. Summary of d 13 C data on TOC n-alkanes and n-fatty acids from the 99 / 1 and 96 / 2 cores from Malham Tarn. Sample 0(96/2) 18(96/2) 25(96/2) 30(96/2) 38(96/2) 365(96/2) 500(96/2) 590(96/2) 235(99/1) 260(99/1) 325(99/1) 360(99/1) 425(99/1) Chara peat 190(99/1) peat 180(99/1)
nC23
nC25
nC27
nC29
n-fatty Acids nC31 nC14 nC16:1 nC16:0 nC18:1 nC18:0 nC22
228.1 231.7 230.3 228.0 230.5 229.1 229.7 232.0 228.9 230.0 229.2 229.9 230.2 232.2 229.6 229.3 230.1 229.2 231.1 229.5 231.3 227.8 221.5 224.2 223.6 229.5 227.2 222.5 228.4 227.7 225.6 238.6 227.7 228.9 228.4 231.5 230.8 227.5 230.8 230.7
231.4 229.8 231.0 231.0 229.5 231.2 230.9 230.4 230.4
232.3 231.8 230.8 233.0 233.4 231.8 231.8 231.1 231.6
232.9 232.9 231.0 232.0 229.6 231.7 232.1 231.5 231.7
233.0 232.9 232.5 232.5 233.6 233.7 233.5 233.2 233.0
225.1 229.2 227.8 229.0 231.2 230.6
227.6 231.4 230.1 230.6 230.4 230.2
229.4 232.3 230.2 231.8 230.7 230.2
231.5 233.9 230.1 232.8
TOC
n-alkanes nC17 nC21
233.1
226.8 227.8 228.1 230.7 230.2 230.8 230.7 228.6 227.2 227.1 224.5
227.1
223.9 233.4 224.6 225.5
226.0
225.1 226.8 225.4
225.3 223.0
nC24
nC26
nC28
226.4 232.2 232.9 226.8 232.5 232.1 232.3 234.3 227.0 233.8 232.9 233.0 228.6 228.8 230.5 226.2 228.0 229.4 236.5 228.9 230.6 232.4 235.7 228.4 230.7 231.7 233.2 229.3 233.2 230.3 231.6 232.9 227.0 233.3 232.7 231.9 224.0 230.4 230.9 230.8 231.0 224.3 221.1 222.1 225.6 225.0 227.4 228.5 222.1 223.6 223.1 224.2 226.5 228.0 234.9 237.7 227.0 231.9 232.1 232.3 232.7 225.7 231.7 231.8 231.7
413 negative in the core. This suggests that Chara is the principal source for these homologues as modern Chara has values of 233.4‰ and 228‰ for n-C 16 and n-C 18 respectively. A recent morphological study of Chara gyrogonites indicates that the Chara in all of 96 / 2 and most of 99 / 1 consists of a single population (authors, unpublished). This evidence supports the possibility of using modern Chara for comparison purposes. In the clay samples, the n-alkanes include short length homologues, and maxima at medium chain length, C 23 and C 24 , except for 325(99 / 1) with a maximum at C 29 . Samples 325(99 / 1) and 425(99 / 1) also show a clear peak of Pristane (Figure 5). This compound was described as a by-product of zooplanktonic digestion of chlorophyll-a (Blumer et al. 1971) pointing to an algal origin, but it is also produced from slow-acting diagenetic reworking of phytol and consequently is a common constituent of hydrocarbons in ancient sedimentary rocks (e.g., Meyers and Ishiwatari (1993b)). It could have been incorporated into the lake sediments by erosion of surrounding rocks. The CPI value, for the n-alkanes of samples 360(99 / 1) and 425(99 / 1) are low, 2.1 and 4.3 respectively (Table 1), indicating mature organic matter. All of these would point the organic matter in clay samples to be a mixture of eroded and autochthonous material of algal origin. No Chara (carbon-
ate) remains have been found in the clay samples but undoubtedly other aquatic macrophytes grew in Malham Tarn (Pigott and Pigott 1959). The lack of detectable C 16 and C 18 n-fatty acids in samples 325(99 / 1) and 360(99 / 1) could be explained by selective diagenetic degradation of these shorter-chain homologues (Meyers and Ishiwatari 1993b). Branched and hydroxy acids were not detected in the sediments, indicating that bacteria did not contribute significantly to the sedimentary organic matter at Malham Tarn. Carbon isotope features Figure 4 shows the profile of d 13 C TOC and biomarkers downcore. For the sake of clarity the nalkanes and n-fatty acids are discussed as 3 groups, (1) long chain length (weighted average of nC 27,29,31 alkanes and n-C 26,28 fatty acids), (2) medium chain length (weighted average of n-C 23,25 alkanes and n-C 24 fatty acids) and (3) short chain length (weighted average n-C 16,18 fatty acids). Short chain homologous for n-alkanes were not detected. The TOC d 13 C depth profile shows three intervals. The basal clay unit and overlying sand / clay / marl unit have d 13 C of 223‰ which decreases at the base of the principal marl unit to a mean value around 230‰, whilst the topmost black marl unit d 13 C increases to
Figure 5. Chromatogram and mass spectra of Pristane in sample 425(99 / 1). Sq.: internal standard.
414 228‰ at the surface. Relatively high d 13 C values in TOC of the last glaciation have been observed from lake sediments in Europe, Africa, Asia and Australia. The study sites range in latitude from Scotland to New Zealand, and in altitude from sea level to over 4200m, and the variation in d 13 C variation is from 1‰ to 27‰ (Street-Perrott et al. 1998). Whittington et al. (1996) found in sediments from a former lake basin, at Lundin Tower, Scotland, a similar behaviour in d 13 C of TOC for the Lateglacial-YD-Holocene transition, similar to that of Malham Tarn. The fact that the d 13 C of TOC is a weighted average of the d 13 C of all organic matter constituents poses a drawback for its interpretation. The long chain n-alkyl lipids represent an allochthonous input from higher plants in the lake catchment. Several factors may have been responsible for the glacial / interglacial change in the carbon isotope composition of higher plants: shifts in the dominant photosynthetic pathway (C3, C4 or CAM), variation in the d 13 C of atmospheric CO 2 , diagenesis and changes in the isotopic discrimination by plants induced by environmental conditions. The C4 / CAM photosynthetic pathways are not relevant here because C4 and CAM plants only occur in areas with a temperate climate (Osmond et al. 1982) and we are considering the Lateglacial at high latitude. Several authors have reported atmospheric CO 2 of d 13 C values lower by 0.3‰ to 1‰ in the last Glacial than in the Holocene (Marino et al. 1992; Leuenberger et al. 1992). These differences are small and in the opposite direction and do not explain our results. A selective diagenetic effect on the isotopic signal could explain the glacial / interglacial change but other authors have reported only minor diagenetic changes in bulk organic matter, around 2‰, (Burchardt and Fritz 1980; Meyers and Eadie 1993; Mayer and Schwark 1999) and even lower changes in n-alkyl lipids (Huang et al. 1997). Meyers and Ishiwatari (1993b) note that in lacustrine sediments having organic carbon concentrations below a few percent, no diagenetic isotope effect occurs. During the Last Glacial, fractionation by C3 plants would have been affected by colder temperatures, dry environment (water stress), lower pCO2 and more open vegetation (reduced canopy effect). Empirical data show that the net effect of these opposing factors on trees was a positive shift of ca. 1.5–4‰ (Krishnamurthy and Epstein 1990; Leavitt and Danzer 1992; Van der Water et al. 1994). This variation would be enough to explain the 3.7‰ maximum difference found in the long chain n-alkanes which are derived from higher plants.
The medium and short chain length n-alkyl lipids i.e, n-C 23 and n-C 25 , medium chain length alkanes, n-C 16 , n-C 18 short chain length fatty acids and n-C 24 medium chain length fatty acids come from autochthonous organic matter, macrophytes and algae, as described above. The factors affecting the d 13 C of these C3 aquatic organisms (Farquhar et al. 1989) are basically the d 13 C and the concentration of dissolved inorganic carbon (DIC) in the lake. DIC d 13 C depends upon its source, i.e, atmospheric CO 2 , carbonates dissolution and organic detritus. The concentration is related with pH, temperature, salinity, aeration and aquatic primary productivity. At this time and at these latitudes tree cover would probably be sparse in the immediate vicinity of the lake. The absence of hazelnuts and pine needles in the clay sediments supports this statement. The results of alkyl lipids homologue distribution (Figure 3) discussed before, point to a mixture of autochthonous and eroded organic matter that would have been unlikely in a well-developed forest. Thus, the possible input of depleted 13 C DIC would be minor. Regarding DIC concentration, clay varving in core 99 / 1 suggest some winter freezing and summer melting that reflects periods of perennial ice cover that will reduce this concentration severely. A different cause for the higher d 13 C values in the biomarker compounds during late glacial may be acidic conditions in the lake water at that time. However as the clay unit contains carbonate in the reactive form of ‘rock flour’ it is unlikely that acidic conditions prevailed, although the pH may have been slightly lower than the current 8.2–8.3. All these factors together would point to a heavy d 13 C and low concentration of DIC, resulting in a heavy d 13 C value of the autochthonous organic matter at Lateglacial interstadial. The change to lighter d 13 C values in the marl, displayed in TOC and in all biomarkers, would correspond to the input of 13 C-depleted DIC from recycled soil organic matter after the post-glacial establishment of vegetation in the catchment. The fact that the d 13 C values of short chain length fatty acids change to lower values more slowly than the others, would represent the time needed for the 13 C depleted inorganic organic carbon to arrive into the lake. In the post-dam event, the shift of d 13 C TOC attributed to the eroded peat, is recorded in the fatty acids, principally, by lighter d 13 C values of medium and long chain homologues and heavier d 13 C of shorter ones. The long chain n-fatty acids, are considered to be from higher plants, but their d 13 C values in the clay section of the core are heavy and more similar to the
415 autochthonous n-alkyl lipid behaviour than the terrestrial higher plants (represented also by the long nalkanes). A possible explanation would be that they are not only coming from terrestrial higher plants but have an important contribution from aquatic macrophytes. Another possible explanation is that the n-fatty acids reflect the same climatic changes more strongly than the n-alkanes due to a different biosynthetic fractionation process. Street-Perrott et al. (1997), Huang et al. (1999) studied the d 13 C of TOC and biomarkers of two African lakes in the Last Glacial–Holocene transition. They found a strong shift in both, TOC and n-alkyl lipids d 13 C, but their results show bigger variations in the n-fatty acids than in the n-alkanes, in agreement with our results. These authors attributed this glacial-interglacial variation to the change in vegetation from C4 to C3, triggered principally by lower CO 2 concentration in the Last Glacial. Very few studies of biomarkers d 13 C have been done in high latitude lakes. Brincat et al. (2000) studied the carbon isotope composition of long chain n-alkanes from Lake Baikal and found no changes throughout the core while the same samples displayed important changes (although smaller than those found in Malham Tarn) in bulk carbon isotope composition initially attributed to changes in vegetation type from C4 to C3. In Malham Tarn both, d 13 C TOC and n-alkyl lipids (long chain n-alkanes and particularly long n-fatty acids), show clear change from Last Glacial to Holocene. It seems that the lower concentrations of CO 2 during the Last Glacial relative to the Holocene would have been a very wide spread phenomenon and would have played a major role in the d 13 C of organic matter. In the high latitude environment of the study area which is devoid of C4 plants, the carbon isotope record of TOC, and biomarker compounds indicative of the allochtonous organic components of the marl lake Malham Tarn, sediments show the effects of changes in pCO2, temperature and humidity in the catchment since the last glacial and allow an assessment of the time lag in the propagation of these changes into the lacustrine environment through the isotope record of the biomarkers of the autochtonous organic matter.
catchment at high latitude since 12Ka. The exclusively C3 flora and the sedimentological and pollen studies enabled a rigorous interpretation of the distribution of TOC, n-alkanes and n-fatty acids and their carbon isotope composition. The transition from Last glacial to Holocene reflected in the change from calcareous sand clay to marl and from pollen assemblages of zones I / II to IV/ V is also marked by a decrease in d 13 C in all the above mentioned organic components. This change agrees with previous reports from study sites spanning a large range of latitude and altitude and seems to be indicative of a widespread-global phenomenon probably due to a major increase in atmospheric CO2 concentration rather than changes in flora resulting from environmental changes that could be local. The decrease in d 13 C is marked most strongly in the long chain and medium chain n-fatty acids and the long chain n-alkanes derived from the terrestrial higher plants in the catchment while the decrease in the short chain fatty acids derived from the lake flora is more gradual covering a span of several thousand years. This lag reflects the time needed for the formation of soil derived CO2 to affect the carbon reservoir of the lake. This phenomenon is not recorded in the d 13 C of the TOC as the TOC is dominated by allochtonous material. The increased erosion of peat from the Tarn Moss since 1791 and changes to the flora of the lake are recorded by an increase in TOC. Again there is a marked difference between the changes in d 13 C of autochtonous and allochtonous biomarker compounds. The detailed record of d 13 C TOC does not show a detectable isotope shift associated with the Little Ice Age. The distribution of biomarkers indicative of Chara are detected in the sediments rich in calcareous remains of Chara thus indicating the importance of its contribution to the carbonate and the biomarker records of marl lakes. This study demonstrates the usefulness of the d 13 C data of individual biomarkers for the elucidation of environmental changes in lakes and their catchment and emphasises the need to distinguish between the respective autochtonous and allochtonous components.
Conclusions The organo geochemical and carbon isotope study of post-glacial sediments from the Malham Tarn (NW England) revealed the evolution of a marl lake and its
Acknowledgements The staff of the Malham Tarn Field Centre are
416 thanked for assistance and laboratory space. P.A Meyers and the anonymous reviewer are thanked for their constructive comments and suggestions. The ´ ˜ at The NERC Isotope Geostay of Rafael Nunez sciences Laboratory, Keyworth was supported by Grant EX99-24200421 of the Spanish Ministry of Education.
References Abrajano T.A. Jr., Murphy D.E., Fang J., Comet P. and Brooks J.M. 1994. 13 C / 12 C ratios in individual fatty acids of marine mytilids with and without bacterial symbionts. Org. Geochem. 21: 611– 617. Barnes M.A. and Barnes W.C. 1978. Organic compounds in lake sediments. In: Lerman A. (ed.), Lakes: Chemistry, Geology, Physics. Springer, Berlin, pp. 127–152. Bird M.I., Summons R.E., Gagan M.K., Roksandic Z., Dowling L., Head J. et al. 1995. Terrestrial vegetation change inferred fromnalkane d 13 C analysis in the marine environment. Geoch. Cosmoch. Acta 59: 2853–2857. Blumer M., Guillard R.R.L. and Chase T. 1971. Hydrocarbons of marine plankton. Mar. Biol. 8: 183–189. Brincat D., Yamada K., Ishiwatari R., Uemura H. and Naraoka H. 2000. Molecular-isotopic stratigraphy of long-chainn-alkanes in Lake Baikal Holocene and glacial age sediments. Org. Geochem. 31: 287–294. Brooks P.W., Eglinton G., Gaskell S.J., McHugh D.J., Maxwell J.R. and Philp R.P. 1976. Lipids in recent sediments I. Straight-chain hydrocarbons and carboxylic acids of some temperature lacustrine and sub-tropical lagoonal / tidal flat sediments. Chem. Geol. 18: 21–38. Burchardt B. and Fritz P. 1980. Environmental isotopes as environmental and climatological indicators. In: Fritz P. and Fontes J.Ch. (eds), Handbook of Environmental Isotope Geochemistry. The Terrestrial Environment Vol. 1. Elsevier, Amsterdam, pp. 473–504. Coletta P., Pentecost A. and Spiro B. 2001. Stable isotope in charophyte incrustrations: relationship with climate and water chemistry. Paleogeogr. Palaeoclim. Palaeoecol 173: 9–19. Collister J.W., Rieley G., Stern B., Eglinton G. and Fry B. 1994. Compound-specific d 13 C analyses of leaf lipids from plants with differing carbon dioxide metabolisms. Org. Geochem. 21: 619– 627. Cranwell P.A. 1974. Monocarboxylic acids in lake sediments: indicators derived from terrestrial and aquatic biota of paleoenvironmetal trophic levels. Chem. Geol. 14: 1–14. Cranwell P.A. 1984. Lipid geochemistry of sediments from Upton Broad, a small productive lake. Org. Geochem. 7: 25–37. Deines P. 1980. The isotopic composition of reduced organic carbon. In: Fritz P. and Fontes J.C. (eds), Handbook of Environmental Geochemistry. The Terrestrial Environment, Part A Vol. 1. Elsevier Scientific Publishing Company, Amsterdam, pp. 329–406. Duan Y., Wen Q., Zheng G., Luo B. and Lanhua M. 1997. Isotopic composition and probable origin of individual fatty acids in modern sediments from Ruoergai Marsh and Nansha Sea, China. Org. Geochem. 27: 583–589.
Eglinton G. and Calvin M. 1967. Chemical fossils. Sci. Am. 216: 32–43. Eglinton G. and Hamilton R.J. 1967. Leaf epicuticular waxes. Science 156: 1322–1335. Farquhar G.D., Ehleringer J.R. and Hubick K.T. 1989. Carbon isotope discrimination and photosynthesis. Annu. Rev. Plant Physiol. Plant Mol. Biol. 40: 503–537. Ficken K.J., Barber K.E. and Eglinton G. 1998. Lipid biomarker, d 13 C and plant macrofossil stratigraphy of a Scottish montane peat bog over the last two millennia. Org. Geochem. 28: 217– 237. Ficken K.J., Li B., Swain D.L. and Eglinton G. 2000. An n-alkane proxy for the sedimentary input of submerged / floating freshwater aquatic macrophytes. Organic Geochemistry 31: 745–749. Freeman K.H., Hayes J.M., Trendel J.M. and Albrecht P. 1990. Evidence from carbon isotope measurements for diverse origin of sedimentary hydrocarbons. Nature 343: 254–256. Fryer G. 1991. Malham Tarn. Biologist 38: 81–83. Giger W., Schaffner C. and Wakeham S.G. 1980. Aliphatic and olefinic hydrocarbons in recent sediments of Greifensee, Switzerland. Geochim. Cosmochim. Acta 44: 119–129. Godwin H. 1975. History of the British Flora. 2nd edn. Cambridge University Press, London, 545 pp. Hakansson S. 1985. A review of various factors influencing the stable carbon isotope ratio of organic lake sediments by the change from glacial to post-glacial environmental conditions. Quat. Sci. Rev. 4: 135–146. Hammarlund D. 1992. A distinct d 13 C decline in organic lake sediments at the Pleistocene-Holocene transition in southern Sweden. Boreas 22: 236–243. Hammarlund D., Edwards T.W.D., Bjorck S., Buchardt B. and Wohlfarth B. 1999. Climate and environment during the Younger Dryas (GS-1) as reflected by composite stable isotope records of lacustrine carbonates at Torreberga, southern Sweden. J. Quat. Sci. 14: 17–28. Holmes P.F. 1965. The natural history of Malham Tarn. Field Studies 2: 199–223. Huang Y., Eglinton G., Ineson P.M., Latter R., Bol D.D. and Harkness D. 1997. Absence of carbon isotope fractionation of individualn-alkanes in a 23 year field decomposition experiment with Calluna vulgaris. Org. Geochem. 26: 497–501. Huang Y., Street-Perrot F.A., Perrott R.A., Metzger P. and Eglinton G. 1999. Glacial-interglacial environmental changes inferred from molecular and compound-specific d 13 C analyses of sediments from Sacred Lake, Mt. Kenya. Geochim. Cosmochim. Acta 63 9: 1383–1404. Jones T.P., Fortier S.M., Pentecost A. and Collinson M.E. 1996. Stable carbon and oxygen isotopic compositions of recent charophyte oosporangia and water from Malham Tarn, U.K.: palaeontological implications. Biogechemistry 34: 99–112. Kolattukudy P.E., Croteau R. and Buckner J.S. 1976. Biochemistry of plant waxes. In: Kolattukudy P.E. (ed.), Chemistry and Biochemistry of Natural Waxes. Elsevier, Amsterdam, pp. 289–347. Krishnamurthy R.V. and Epstein S. 1990. Glacial-interglacial excursion in the concentration of atmospheric CO 2 : effect in the 13 C / 12 C ratio in wood cellulose. Tellus 42b: 423–434. Leavitt S.W. and Danzer S.R. 1992. 13 C variations in C 3 plants over the past 50,000 years. Radiocarbon 34: 783–791. Leuenberger M., Siegenthaler U. and Langway C.C. 1992. Carbons isotope composition of atmospheric CO 2 during the last ice age from an Antartic ice core. Nature 357: 488–490.
417 Lockheart M.J.,Van Bergen P.F. and Evershed R.P. 1997.Variations in the stable carbon isotope compositions of individual lipids of modern angiosperms; implications for the study of higher land plant-derived sedimentary organic matter. Org. Geochem 26: 137–153. Marino B.D., McElory M.B., Salawich R.J. and Spaulding W.G. 1992. Glacial-to-interglacial variations in the carbon isotopic composition of atmospheric CO 2 . Nature 357: 461–466. Mayer B. and Schwark L. 1999. A 15,000-year stable isotope record from sediments of Lake Steisslingen, Southwest Germany. Chem. Geol. 161: 315–337. Meyers P.A. and Eadie B.J. 1993. Sources, degradation and recycling of organic matter associated with sinking particles in Lake Michigan. Org. Geochem. 20: 47–56. Meyers P.A. and Ishiwatari R. 1993a. Lacustrine organic geochemistry. An overview indicators of organic matter sources and diagenesis in lake sediments. Org. Geochem. 20: 867–900. Meyers P.A. and Ishiwatari R. 1993b. The early diagenesis of organic matter in lacustrine sediments. In: Engel M.H. and Macko S.A. (eds), Organic Geochemistry. Principles and Applications. Plenum Press, New York, pp. 185–209. ¨ ¨ Oldenburg T.B.P., Rullkotter J., Bottcher M.E. and Nissenbaum A. 2000. Molecular and isotopic characterization of organic matter in recent and sub-recent sediments from the Dead Sea. Org. Geochem. 31: 251–265. O’Leary M.H. 1981. Carbon isotopic fractionation in plants. Phytochemistry 20: 553–567. Osmond C.B., Winter K. and Ziegler H. 1982. Functional significance of different pathways of CO 2 fixation in photosynthesis. In: Lange O.L., Nobel P.S., Osmond C.B. and Ziegler H. (eds), Physiological Plant Ecology II. Water relations and carbon assimilation. Springer-Verlag, New York, pp. 479–547. Ostrom P.H., Ostrom N.E., Henry J., Eadie B.J., Meyers P.A. and Robbins J.A. 1998. Changes in the trophic state of Lake Erie: discordance between molecular d 13 C and bulk d 13 C sedimentary records. Chem. Geol. 152: 163–179. Pentecost A. 2000. Some observations on the erosion of Tarn Moss by the waters of Malham Tarn. Field Studies 9: 569–581. ˜ R. 2000. PalaeoenPentecost A., Spiro B., Coletta P. and Nunez vironmental interpretation of the early Postglacial sedimentary record of a marl lake using biostratigraphy and stable isotopes (Malham Tarn, North Yorkshire) Conf. Abs.5 (2000):86. Cambridge Publications (eds).
Pigott C.D. and Pigott M.E. 1963. Late-glacial and post-glacial deposits at Malham, Yorkshire. New Phytologist 62: 317–343. Pigott M.E. and Pigott C.D. 1959. Stratigraphy and pollen analysis of Malham Tarn and Tarn Moss. Field Studies 1: 84–101. Rieley G., Collier R.J., Jones D.M., Eglinton G., Eakin P.A. and Fallick A.E. 1991. Sources of sedimentary lipids deduced from stable carbon-isotope analyses of individual compounds. Nature 352: 425–426. Sargent J.R. and Whittle K.J. 1981. Lipids and hydrocarbons in the marine food web. In: Longhurst A.R. (ed.), Analysis of Marine Ecosystems. Academic Press, London, pp. 491–533. Shameel M. 1990. Phycochemical studies on fatty acids from certain seaweeds. Bot. mar. 33: 429–432. Spooner N., Rieley G., Collister J.W., Lander M., Cranwell P.A. and Maxwell J.R. 1994. Stable carbon isotopic correlation of individual biolipids in aquatic organisms and a lake bottom sediment. Org. Geochem. 21: 823–827. Street-Perrott F.A., Huang Y., Perrott A., Eglinton G., Baker P., Khelifa L. et al. 1997. Impact of lower atmospheric carbon dioxide on tropical mountain ecosystems. Science 278: 1422– 1426. Street-Perrott F.A., Huang Y., Perrott R.A. and Eglinton G. 1998. Carbon isotopes in lake sediments and peats of last glacial age; implications for the global carbon cycle. In: Griffiths H. (ed.), Stable Isotopes Integration of Biological, Ecological and Geochemical Process. Bios Scientific Publishers Ltd, Oxford, pp. 381–397. Turney C.S.M., Beerling D.J., Harkeness D.D., Lowe J.J. and Scott E.M. 1997. Stable carbon isotope variations in northwest Europe during the last glacial-interglacial transition. J. Quat. Sci. 12: 339–344. Van der Water P.K., Leavitt S.W. and Betancourt J.L. 1994. Trends in stomatal density and 13 C / 12 C ratios of Pinus flexilis needles during the last glacial-interglacial cycle. Science 264: 239–243. Viso A.C., Pesando D., Bernard P. and Marty J.C. 1993. Lipid components of the Mediterranean seagrass Posidonia Oceanica. Phytochemistry 34: 381–387. Whittington G., Fallick A.E. and Edwards K.J. 1996. Stable oxygen isotope and pollen records from eastern Scotland and a consideration of Late-glacial and early Holocene climate change for Europe. J. Quat. Sci. 11: 327–340.
