Earth and Planetary Science Letters 233 (2005) 17–32 www.elsevier.com/locate/epsl
In situ measurement of seasonal d18O variations and analysis of isotopic trends in a modern speleothem from southwest Australia
P.C. Treblea,b,T, John Chappellb, Michael K. Gaganb, K.D. McKeegana, T. Mark Harrisona,b
aDepartment of Earth and Space Sciences, University of California, Los Angeles, CA 90095-1567, USA bResearch School of Earth Sciences, The Australian National University, Canberra ACT 0200, Australia Received 27 September 2004; received in revised form 20 January 2005; accepted 17 February 2005 Available online 29 March 2005 Editor: E. Bard
Abstract
We present a record of seasonal and inter-annual oxygen (d18O) and carbon (d13C) isotope ratios from an 81-year-old stalagmite from Moondyne Cave, southwest Australia. The growth history of stalagmite MND-S1 is known since it grew on a cave boardwalk that was installed in 1911 and removed in 1992. This stalagmite provides an excellent test of speleothem climate proxies because the regional climate is strongly seasonal (wet winter/dry summer) and has experienced a 200 mm (20%) reduction of mean rainfall since the mid-1960s, and a 0.8 8C temperature rise since ~1953. Seasonal variations in calcite d18O were measured in situ by high spatial resolution ion microprobe, whilst inter-annual variations of d18O and d13C were measured by conventional gas-source mass spectrometry. Comparison of the speleothem stable isotopes and instrumental temperature records reveals that d18O variations are too large to be driven by temperature alone, and are in the opposite sense. However, daily rainfall d18O measurements show that the mean seasonal range in d18O of rainfall in southwest Australia is large (2x) and inversely correlated with rainfall amount. A rainfall driver for the speleothem d18O is confirmed by the detection of seasonal shifts of 0.7–1.5x in speleothem d18O that track rainfall d18O, smoothed by storage in the overlying limestone. The seasonal range in speleothem d18O is larger than any inter- annual and decadal variation observed in the record. The prominent annual cycles in speleothem d18O revealed by ion microprobe analysis indicate that subtle changes in the frequency of intense winter rainfall events, or possibly also moisture sources, could produce significant changes in mean speleothem d18O. The ion microprobe results also raise the possibility that the masses of speleothem calcite deposited in winter and summer could vary as a function of the seasonal drip rate and carbonate saturation state of these waters. If this is the case, then small changes in the relative masses of calcite deposited in winter and summer could produce significant shifts in mean d18O and d13C that have a complex relation to climate. This finding should be generally applicable to the interpretation of long-term trends in speleothem geochemical records for shallow cave
T Corresponding author. Present address: Research School of Earth Sciences, The Australian National University, Canberra ACT 0200, Australia. Tel.: +61 2 6125 3406; fax: +61 2 6125 0738. E-mail address: [email protected] (P.C. Treble).
0012-821X/$ - see front matter D 2005 Elsevier B.V. All rights reserved. doi:10.1016/j.epsl.2005.02.013 18 P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 sites where seasonal variations in geochemical tracers are relatively large, including most of the sub-tropical monsoon belts and mid to high latitudes with distinct rainfall seasons. D 2005 Elsevier B.V. All rights reserved.
Keywords: speleothem; oxygen isotopes; carbon isotopes; annual cycle; rainfall; ion microprobe
1. Introduction raise the 13C/12C of the precipitating solution. This may be driven by a reduction in the partial pressure of 13 The stable isotopes of carbon (d C) and oxygen CO2 in drip waters when they enter the cave [11,16], (d18O) in speleothem carbonates are widely regarded or because calcite has been lost from solution in as climate proxies and are routinely measured [1–7], aquifer pore spaces or over other formations (prior yet few modern records exist to test their veracity calcite precipitation) [12]. against the instrumental climate record. This step is A critical factor in interpreting speleothem d13C critical for understanding climate proxies, particularly and d18O records is determining whether disequili- those reflecting cumulative signals of complex pro- brium isotopic fractionation caused by rapid CO2 cesses. Speleothem d18O records the combined degassing or, rarely, by evaporation of the precipitat- signatures of rainfall d18O and cave temperature, ing solution [11], overprints climatic signals. Isotopic provided the calcite is precipitated in oxygen isotopic disequilibrium is indicated by a simultaneous increase equilibrium [1]. Methods for separating the two of d13C and d18O down a growth layer from the effects, such as the analysis of fluid inclusions that speleothem apex, and by correlation between d13C represent the original drip water [8], are complicated and d18O along a growth axis [11]. This is because and not routinely performed. Thus, many speleothem d13Candd18O are affected simultaneously if the records are interpreted as combined records of rainfall removal of CO2(g) outpaces the conversion of CO2(aq) and temperature [4,5], while some are recognised as to HCO3 [11]. However, these tests tend not to be being clearly dominated by rainfall [6]. definitive in practice because it is almost impossible to Speleothem d13C is also a composite signal, sample coeval calcite along a growth layer. Moreover, reflecting biogenic fractionation of the soil carbon climatic factors conceivably can simultaneously affect pool [9], limestone dissolution under open- or closed- axial profiles of both speleothem d18O and d13C. system conditions [10,11], prior carbonate precipita- To improve our understanding of isotopes in tion from groundwater [12], and the rate of CO2 speleothems, we examined seasonal variations of degassing from drip waters over a speleothem surface d18O, and inter-annual variations of d18Oandd13C [10,11]. The term open system or closed system refers in a modern speleothem of known growth history. Our to limestone dissolution conditions where ground- test sample is a 33 mm tall, 59 mm wide stalagmite waters either remain in contact with soil CO2 (open) (MND-S1) that grew from 1911 to 1992 on a board- or are isolated from soil CO2 (closed) after acid- walk in Moondyne Cave, southwest Western Australia ification [11]. Biogenic CO2 produced under C3 (34816VS, 115805VE; Fig. 1). The calcite crystal fabric vegetation has an average d13C value of 27x, but of MND-S1 is columnar palisade [17], with inclusions may vary by approximately F5x depending on the between crystallites preserved in the mid-section environmental effects on photosynthesis [13,14]. Soil (open columnar palisade fabric; 1930–1970), and CO2 is dominated by biogenic CO2 but can be mixed dense, compact inclusion-free (closed columnar pal- with atmospheric CO2 when soil respiration rates are isade fabric) calcite above and below this (Fig. 2). low [15]. The contribution of carbon atoms from Stalagmite MND-S1 preserves clear trace element limestone dissolution is negligible under open-system cycles (P, Mg, U, Sr, Ba, and Na) previously analysed conditions and small under closed-system conditions by excimer laser ablation inductively coupled plasma [11,16]. By contrast, degassing of CO2 before drip mass spectrometry (ELA-ICPMS) [18]. These cycles water reaches the speleothem, and the subsequent were determined to be annual based on the agreement conversion of HCO3 to CO2(aq), may substantially between the number of cycles in MND-S1 and the P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 19
Fig. 1. Location of stalagmite MND-S1 (Moondyne Cave), rainfall collection site (Cape Leeuwin), Cape Naturaliste, and high-quality instrumental climate dataset (Manjimup) in southwest Western Australia. A long and detailed study of rainfall d18O was conducted at Hobart, Tasmania. Area south of thick grey line is Australia’s winter rainfall zone. number of years that the boardwalk was in the cave also. We used high-resolution in situ ion microprobe [18]. Thus the chronology of MND-S1 is well-known, analysis to investigate this possibility, and hence that constrained by the time the boardwalk was in place rainfall is the primary driver of d18O in MND-S1. (1911–1992) and the yearly markers provided by the annual trace element cycles. The trace element cycles represent seasonal variations in the chemistry of 2. Environmental setting meteoric water as it passes through, and interacts with, the soil and limestone. No detailed study of Moondyne Cave is a syngenetic karst cavern aquifer hydrology was undertaken, but the presence of developed in Late Cenozoic calcarenite and aeolianite seasonal trace element variations led us to expect that limestone [19], with a relatively thin roof (15–20 m) seasonal variations in rainwater d18O will be preserved beneath a Eucalyptus forest. Total cave length is 170
Fig. 2. Thin-section photograph of stalagmite MND-S1 under cross-polarised light showing the transition from compact, inclusion-free columnar fabric to inclusion-rich columnar fabric at ~1930. The photograph spans ~1917 to ~1940. The abrupt change in growth fabric coincides with the 0.3x increase in speleothem d18O at 1930. 20 P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 m and MND-S1 was located approximately 60 m mental records are needed to understand the nature of from the cave entrance (approximately 1.5 1.5 m this climatic change. Hence, it is important to fully wide). The cave is located 3 km from the coast in a evaluate the potential for reconstructing rainfall temperate climate with predominantly winter rainfall. variations from speleothem oxygen and carbon iso- Average annual rainfall at the nearest station, Cape tope records. Leeuwin Lighthouse (12 km from Moondyne Cave; The oxygen isotopic composition of rainfall is Fig. 1), is 1010 mm. 798 mm/year is recorded at known to vary both in time and space [22] and can Manjimup station (97 km inland from Moondyne strongly influence speleothem d18O when the sea- Cave; Fig. 1) where the nearest high quality climate sonal variation is large [6]. To determine the seasonal dataset is available [20]. Average annual surface air variation of d18O in southwest Australian rainfall, we temperatures at Cape Leeuwin range from 13.8 8Cin also present a 1 year (October 2000–October 2001) July–August to 20.3 8C in January–February, which is time series of rainfall d18O using samples collected on much larger than the annual temperature range where a daily basis at Cape Leeuwin Lighthouse. This MND-S1 grew (15.9–16.5F0.2 8C), recorded by data collection would have been ideally complemented logger measurements at 4-h intervals for 1 year. The by weekly or monthly cave drip water sampling over 0.6 8C seasonal cycle in cave temperature lags the the same time period; however, it was not possible to surface air temperature by approximately 3 months, collect drip water, apart from on one occasion in June owing to the insulation provided by the limestone. 2002, owing to cave access restrictions. The average daily temperature range at the lighthouse is approximately 5–6 8C, but no diurnal temperature variations were recorded in the cave, indicating 3. Methods minimal air circulation at the site of MND-S1. The area above Moondyne Cave was reasonably 3.1. d18O analysis of rainfall free of activities that would have greatly disturbed the soil and catchment owing to the hillocky topography Rainfall samples were caught in a clean high- of the limestone ridge, although the area was regularly density polyethelene bottle fitted with a funnel and burnt to reduce the understorey of the Eucalyptus ball valve and collected between 8 and 9 a.m. after forest to allow cattle passage. The forest understorey rain fell during the previous 24 h, between October may have thickened through time in response to less 20, 2000 and October 12, 2001, except on four frequent burning when large tracts of land, including occasions when the rain was collected after 48 h. d18O that over Moondyne Cave, were incorporated into the measurements were conducted on samples for which Leewin–Naturaliste National Park in about 1980 (S. at least 1 mm of rain fell. The d18O values of rainfall Eberhard, personal communication). Moondyne Cave collected between October 2000 and May 2001 were was used for cave tours until 1959 and was closed measured in duplicate at the Research School of until 1992 when it underwent restoration, and the Biological Sciences, ANU, using an on-line pyrolysis boardwalk on which MND-S1 grew was removed. No technique [23]. Twenty-six measurements of the in- detailed records of tourist numbers exist between 1911 house ANU-HP1 water standard ( 5.54xVSMOW) and 1959 when tours were operated, but we believe yielded a precision of F0.5x (2r). The d18O values that tours were infrequent as tourists had to first of rainfall collected between June 2001 and October collect the guide from Augusta, 10 km from the cave. 2001 were measured off-line at the Research School Regional meteorologic records show that mean of Earth Sciences (RSES), ANU, following a modi- winter temperature has risen by 0.8 8C since 1954, fied version of the H2O–CO2 equilibration technique and rainfall has decreased by about 200 mm (20%) [24] described by Socki et al. [25]. Two-millilitre since 1965–1970 [21]. This has resulted in a 42% water samples were equilibrated with 8 ml of CO2 gas decrease of streamflow into water storages. Conse- of known isotopic composition at 25.0F0.1 8C for a quently, the city of Perth and the surrounding region minimum of 40 h. The equilibrated H2O–CO2 gas was face a critical shortfall of water. Palaeo-rainfall then injected into a vacuum line for CO2 extraction. 18 records that extend climate data beyond the instru- The d O value of the purified CO2 was measured on P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 21 a Finnigan MAT-251 mass spectrometer and enough to permit sub-annual sampling by micro- expressed relative to Vienna Standard Mean Ocean shaving techniques [28] (e.g., every 30 Am along the Water (VSMOW). Nine measurements of ANU-HP1 growth axis). However, to obtain the 200 Agof yielded a precision of F0.04x (2r). Measurements powder required for analysis, the speleothem must be of two unknowns using both the pyrolysis and CO2 shaved deeper along the layers, perpendicular to the equilibration techniques were within 0.5x. growth axis (1.4 1.4 mm). This technique is only The rainfall d18O data were compared with daily suitable if the growth layers are parallel to avoid rainfall totals and mean, maximum and minimum smoothing the signal by sample averaging. We surface temperatures recorded during each 24-h observed from secondary electron microscopy images collection period at the lighthouse. Partial correlation [29] that the annual growth layers in MND-S1 are not coefficients are presented to determine the relation- simple parallel structures, but vary in thickness and ship between d18O and rainfall or temperature with occasionally pinch-out. This complex growth pre- the influence of the inter-related rainfall and temper- cludes the accurate analysis of annual cycles by the ature variables removed (r = 0.22, p b0.05, for the micro-shaving techniques described. We also point- raindays examined). out that 100% recovery of finely powdered material is difficult, and much smaller amounts would actually be 3.2. Analysis of speleothem d18O and d13Cby recovered (50–100 Ag), reducing the analytical conventional techniques precision. As the micro-shaving techniques would average the MND-S1 d18O signal, we employed high A 5 mm thick slice was cut from the central growth spatial resolution in situ ion microprobe techniques to axis of stalagmite MND-S1 and then halved along the measure sub-annual d18O. Sputter craters from ion growth axis. Carbonate powders were milled in a microprobe techniques are ~20–30 Am in diameter continuous fashion along the cut edge of the growth and 1–2 Am deep. axis using a 2 mm diameter tungsten carbide end-mill In situ d18O was measured on a Cameca IMS-1270 bit, effectively shaving approximately 1 mg of powder ion microprobe at the University of California, Los at 0.5 mm increments. This sampling technique fully Angeles. Polished Au-coated samples were sputtered integrates the isotopic signal for every 0.5 mm with a 2–3 nA Cs+ primary beam to produce negative increment of growth, as opposed to the more common secondary ions at mean intensities of approximately method of sampling speleothems by drilling isolated 3 109 cps (16O) and 6 106 cps (18O). The 18O and holes. The technique adopted here is preferable when 16O ion beams were collected simultaneously in potentially large d18Ovariations over short intervals, Faraday cups after pre-sputtering the analytical area such as seasonal variations, are anticipated. Full for 180 s to achieve sputtering equilibrium. Signals details of the method are given in Gagan et al. [26]. were measured for 100 s (collected in 10 blocks of 10 200-Ag sub-samples were reacted with anhydrous s integration time), which yielded a typical internal 16 18 H3PO4 in a Kiel microcarbonate preparation device at counting precision of F0.1x (2r). The O and O 90 8C, equivalent to the McCrea method [27]. The intensities were maintained within a narrow range to evolved CO2 gas was measured in a Finnigan MAT avoid potential need to correct for non-linear detector 251 mass spectrometer at RSES, ANU. Data are response. Background current measurements were normalised to the Vienna Peedee Belemnite (VPDB) made for every 10 unknowns by blocking the Cs+ scale using National Bureau of Standards NBS-19 beam and subtracting from sample measurements. (d18O= 2.20x and d13C=+1.95x). The long-term The analytical precision of ion microprobe meas- measurement precision for NBS-19 at RSES is urements of d18O in individual grains was F0.7x F0.07x (2r) for d18O and F0.04x (2r) for d13C. (2r, n =14) for NBS-19 and F0.5x (2r, n =8) for NBS-18. In principle, this level of uncertainty could 3.3. Ion microprobe analysis of seasonal d18O reflect the inherent instability of the ion microprobe analyses with respect to reproducibility of instrumen- Annual cycles of trace elements in MND-S1 have tal mass discrimination. However, isotopic hetero- an average length of ~370 Am [18], which is wide geneities of NBS-19 and NBS-18 at the spatial scale 22 P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 of ion microprobe analyses may also contribute to the imise spatial resolution across the fine-scale growth observed variability. To test this idea, we performed bands. Sputter ablation pits were kept ~30 Am apart closely spaced in situ d18O analyses on a petrographic to avoid undesirable topographic effects on the glass slide. The results showed that the ion microp- trajectories of charge-neutralising electrons, as well robe can produce repeat measurements with a as Cs+ and O ions. The positioning of each ion precision of F0.4x (2r,n =7), once stable running probe sputter crater was determined using a Mitutoyo conditions are achieved, on a (presumably) homoge- profile projector with the digital readout on the x–y neous sample with a near-perfectly flat geometry. stage reproducible to F2 Am. The relationship Regardless of whether the variations in the between d18O and trace element variations was analyses of NBS-19 or NBS-18 can be attributed to established by measuring trace elements by ELA- isotopic heterogeneity, or geometric imperfections ICPMS, effectively scanning directly across the caused by the mounting and polishing of small transects of in situ d18O pits with a 32 Am spot as NBS-19 and NBS-18 grains in epoxy, we considered described previously [18]. This technique produced these to be unsuitable standard materials for calibra- trace element concentration data measured over the tion of unknowns with small isotopic variations, such same transect and thus is accurately registered as speleothems. The isotopically homogeneous petro- spatially with the ion microprobe d18O data. In order graphic glass slide is useful for monitoring potential to compare the conventional and in situ isotope data, measurement drift but cannot be used to standardise an age model was constructed for the conventional speleothem d18O owing to possible matrix effects d18O record using the annual Ba cycles measured by between dissimilar materials. Therefore, the means of ELA-ICPMS along the cut edge that was created all in situ measurements of d18O in each transect were when the conventional samples were milled. corrected to the bulk mean d18O value determined from conventional analyses (e.g., [30]). While the accuracy of d18O values (relative to VPDB or 4. Results and discussion VSMOW) is of secondary importance for our inves- tigation, this standardisation procedure effectively 4.1. Rainfall d18O improves measurement reproducibility [30],thus permitting resolution of isotopic variations intrinsic Strong seasonality in rainfall d18O is evident in the to the sample. The difference of 21.8F0.9x in d18O Cape Leeuwin Lighthouse record (Fig. 4A) with between NBS-19 and NBS-18, measured by the precipitation-weighted d18O lower in wet winter UCLA ion microprobe, is close to the accepted dif- months ( 4.7x, May–September; October incom- ference of 20.8x, so no adjustment to the relative pletely sampled) and higher in dry summer months instrumental mass discrimination was necessary. ( 2.7x, November–April). Overall, the precipita- Analytical spots were arranged in a zig-zag pattern tion-weighted mean d18O value is 4.3x. Fig. 4B (Fig. 3) along the speleothem growth axis to max- shows rainfall and d18O data grouped by months with
Fig. 3. Arrangement of ion microprobe ablation pits on stalagmite MND-S1. Individual measurements were positioned in a zig-zag pattern to increase the spatial resolution of the record. Arrows indicate the direction of measurements. Trace element variations were later measured directly over the ion microprobe pits by laser ablation along each of transects 1 and 2. Transect 3 (not shown) is located 1.5 mm from, and parallel to, transects 1 and 2. P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 23
) 6 A 1/2 4
2 Rainfall (mm
0 ) 0 -2 VSMOW -4
O (‰ -6 18
δ -8 11/1/200012/1/2000 1/1/2001 2/1/2001 3/1/2001 4/1/2001 5/1/2001 6/1/2001 7/1/2001 8/1/2001 9/1/2001 10/1/2001 Date (mm/d/yyyy) y = -0.64x - 1.25 Rainfall (%) 20 B 0 2 r = 0.44 C 15 ) 10 -2 5 -4 0 VSMOW 25 -6 20 15 O (‰ -8 18
10 δ 5 -10
O signal (%) 0
18 1 2 3 4 5 6 δ O N D J F M A M J J A S O 2000 2001 Rainfall (mm1/2) Month
Fig. 4. Comparison of the square root of rainfall amount and d18O of rainfall in southwest Australia. (A) Daily rainfall d18O measurements at lighthouse from October 20, 2000 to October 12, 2001. The mean d18O of winter rainfall (May–October) is 2x lower than the that for summer rainfall (November–April). Dashed lines show rainfall events associated with meridional troughs identified from mean sea-level pressure and cloud cover synoptic charts. (B) Data grouped into months and expressedP as a portion of the yearly rainfall and d18O values demonstrating the 18 dominant winter signal. Monthly d18O expressed as a percentage is ðÞd Oi rainfalli mean d18O 100, where i =events total rainfall collected precipitation weighted i1, i2, i3,. . ., in per month and n =number of events per month. (C) Cross-plot showing non-linear inverse relationship between rainfall amount and rainfall d18O. both variables expressed as the percentage each month An inverse relationship with rainfall amount and contributes to the overall signal in the October 2000– no significant relationship between rainfall d18O and October 2001 dataset. This exercise further reveals the surface temperature is consistent with the findings for strong seasonal rainfall d18O cycle and demonstrates a 5-year event-based rainfall d18O record [31] from that the majority of drip water d18O will be derived Hobart, Tasmania, in southeastern Australia (location: from May to September rainfall. Fig. 1). Both sites lie in the path of the same mid- The partial correlation coefficient between d18O latitude low-pressure systems and fronts that first and rainfall amount is higher (r = 0.60; p b0.01) cross or pass south of southwest Australia, then cross than that for mean, maximum, or minimum surface Tasmania a few days later. In Tasmania, rainfall d18O temperatures (r =0.03–0.14) based on the daily data. varies inversely with rainfall amount, and both The relationship between rainfall amount and d18Ois variables are closely related to the prevailing synoptic inverse and non-linear over the large range of daily conditions, driven by the intensity and proximity of rainfall sampled (1–38 mm), with heavier rainfall low pressure systems in the Southern Ocean [31]. events (10 mm or more) being relatively isotopically Large amounts of isotopically lighter rain are deliv- enriched. If the square root of rainfall amount is taken ered when these systems pass close to the coast of (Fig. 4C), the partial correlation coefficient between southern Australia, whereas lesser amounts of isotopi- that parameter and rainfall d18Ois 0.66 ( p b0.01). cally heavier rain are delivered from weaker frontal 24 P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 rainfall generated when these systems pass far to the from about 2.7x to 4.8x. The long-term trend south of Australia [31]. This trend is driven by the in the ion microprobe data is 1x compared with intense convective and recycling processes in these 0.2x for the conventionally measured data. This cyclonic systems such that the 18O/16O ratio is lowest discrepancy cannot be due to temporal analytical at the core compared with frontal rainfall produced drift as tracks 1 and 2 were acquired in opposite further from the centre, where atmospheric convection directions (Fig. 3), which suggests that a systematic is less intense [32]. error due to sample geometry was responsible. Subtle relief between the speleothem and surround- 4.2. In situ ion microprobe d18O: seasonal climatic ing epoxy resin, due to polishing of dissimilar signals materials, can affect instrumental mass fractionation by altering secondary ion optics and the charge Two parallel transects of in situ ion microprobe neutralisation balance at the sample edge [33]. The measurements of d18O from a 2 mm section of trend in the microprobe d18O record was corrected to MND-S1 (~45 measurements per track) equivalent to 0.2x by subtracting an extrapolated linear correction ~6 yr growth are shown in Fig. 5A, and conven- for each transect (Fig. 5C). tional d18O analyses from the same interval in Fig. Although somewhat noisy, the patterns in d18O 5B. The four conventional analyses vary by only for both microprobe tracks are similar. The F0.4x 0.2x, whereas d18O for the microprobe tracks varies (2r) error bar for the microprobe d18O data in Fig. 5A and C is based on the measurement reproduci- bility for a glass slide and demonstrates that some of 18 -2.5 A the high frequency d O variability in MND-S1 is within analytical uncertainty. However, the three -3.0 d18O cycles between 3.5 and 4.4 mm in Fig. 5C -3.5 are clearly reproduced in both microprobe transects. 18 -4.0 Another d O cycle between 4.5 and 4.7 mm is reproduced, but appears to be somewhat asynchro- -4.5 18 conventional δ O nous. The wavelengths of these cycles (0.3–0.4 mm) ) -5.0 ion microprobe track 1 are similar to the average annual trace element
VPDB ion microprobe track 2 wavelength (0.37 mm).
‰ 18 -3.5 B Fig. 6 shows ion microprobe d O data for a third, O (
18 longer transect located further from the sample edge,
δ -4.0 1.5 mm away from and parallel to transects 1 and 2. -2.5 18 C The ion microprobe and conventional d O records -3.0 are in agreement, with neither showing any overall -3.5 trend, indicating that transect 3 was not affected by geometric effects. Transect 3 d18O data are shown in -4.0 parallel with Ba and Mg, which vary seasonally with -4.5 low Ba and high Mg in summer [18]. Judged from the Ba master record constructed for MND-S1 [18], -5.0 transect 3 spans years 1975–1985 AD. Variations of 4.8 4.4 4.0 3.6 3.2 2.8 microprobe d18OinFig. 6 can be matched visually to Distance from top (mm) trace element cycles at several points, but agreement Fig. 5. Comparison of ion microprobe d18O and conventional d18O is weaker than that between the trace element cycles records for stalagmite MND-S1. (A) Ion microprobe d18O themselves. Where agreement is clear, d18O maxima demonstrating reproducibility for two adjacent transects on MND- 18 correspond to high Mg and low Ba in summer, S1. (B) Conventional d O record is shown for comparison. (C) Ion 18 microprobe d18O after correcting for geometric error. These data whereas d O minima correspond to low Mg and show three to four reproducible cycles between 3.5 mm and 4.7 high Ba in winter (e.g., years 1979 and 1980). The 18 mm. Bar indicates 2j uncertainty. amplitudes of the most regular d O cycles are ~0.7– P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 25
1979 1980 14 1976 1977 1982 1975 1978 1984 12 1981 1983 1985 10
Ba (ppm) 8 4500 4000 3500 3000 Mg (ppm) 2500 )
-3.0 drier VPDB
‰ -3.5
O ( -4.0 18 δ wetter
4.6 4.2 3.8 3.4 3.0 2.6 2.2 1.8 1.4 1.0 Distance from top (mm)
Fig. 6. Comparison of a third ion microprobe d18O transect (circles), conventional d18O record (squares), and annual cycles of Ba and Mg measured by ELA-ICPMS [18] over the same section on MND-S1 (~1975–1985). d18O peaks correspond with Mg peaks and Ba troughs (dashed lines) when prominent annual cycles are present in the signals. Bar indicates 2j uncertainty for microprobe d18O.
1.5x. However, some years show more than one Australia [35], presumably because seepage-fed com- d18O peak (e.g., 1981 and 1982), and major d18O ponents dominate over fracture-fed components (e.g., maxima do not always coincide with Ba minima [36]), resulting in smoothing of the rainfall d18O (e.g., 1977). cycle. The 0.7–1.5x cycles in speleothem d18O are We consider that a major source of the speleo- consistent with the range expected for a smoothed them d18O variation lies in seasonal variation of the seasonal cycle of rainfall d18O. Finally, it is not isotopic composition of rainfall, driven primarily by unexpected that trace element cycles, which essen- rainfall amount and not by surface temperature. tially follow the seasonal cycle of rainfall [18], are Given that the observed seasonal temperature more regular than d18O variations, which reflect the variation in Moondyne Cave is only 0.6 8C, the complex mixture of the isotopic signatures of indi- seasonal difference in the temperature-dependent vidual rainfall events. This is a noisy process. Fig. 4A fractionation of 18O/16O in speleothem calcite will shows that event-to-event d18O values can vary by as be negligible (~0.14x based on 0.23x/8C; [34]). much as 6–7x. Moreover, if the temperature-dependent fractionation effect was dominant, speleothem calcite d18O would 4.3. Inter-annual speleothem d18O and d13C: complex tend to be lower in summer than in winter, which is responses to climate opposite to the observed pattern. Thus the tendency for d18O in MND-S1 to be lower in winter is The inter-annual d18O and d13C data are shown consistent with the relationship between rainfall alongside the instrumental temperature and rainfall amount (wetter in winter) and rainfall d18O (lower records in Fig. 7. The d13C values are relatively in winter) deduced from measurements of d18Oin constant at 9x between 1912 and 1925, and local rainfall (Fig. 4A–C). decrease to 10x by 1935. d13C values increase The seasonal range of d18O in Moondyne Cave towards the present, with minor lows at 1950, 1965, drip water is not known. However, it is likely to be and 1974, to reach 7.7x in 1992. Speleothem smaller than the 2x seasonal range in rainfall d18O, d18O values began at 3.8x in 1912, then increased as is observed for shallow caves elsewhere in southern to 3.5x in 1992 and showed several inter-annual 26 P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32
Year 1920 1930 1940 1950 1960 1970 1980 1990 Ann. temperature
2 anomal 1 0 ° -1 +0.8 C -2 y
-3
6 rise Rainfall decrease 4 Temperature -200 mm 2
0 δ 13 May-October
-7.5 C ( rainfall anomaly -8.0
-8.5 ‰
-9.0 VPDB
) -9.5 -10.0 -3.4 ) VPDB -3.5 ‰ -3.6
-3.7 Mg anomaly O ( -3.8 18 2 δ 1 0 -1 800 -2 600 m/a)
µ 400 ( 200 Growth rate 1920 1930 1940 1950 1960 1970 1980 1990 Approximate year
Fig. 7. Comparison of annual temperature, May–October rainfall, conventional d13C and d18O, speleothem Mg [18], and speleothem growth rate for MND-S1. Black curves show the three-point running means. Instrumental temperature, rainfall (Manjimup dataset [20]), and Mg anomalies are calculated by subtracting the mean and dividing by the standard deviation of the period 1965–1992 (units are standard deviation). Annual average temperature has increased by 0.8 8C since 1953 while winter rainfall has decreased by 200 mm since 1964 in the Manjimup record. Growth rate is shown as the yearly wavelength (dashed curve) and 3-yr running mean (solid curve) of a master trace element record compiled from 11 ELA-ICPMS transects on MND-S1 [18]. The growth rates differ slightly from those published previously [18] as this curve was determined using the distance over successive 5-yr intervals. variations that mimic details of the d13C record. lower during this interval than at any time before or However, there is a significant positive shift of 0.3– since. 0.5x between 1930 and 1955 that is not mirrored in Neither d13C nor d18O shows simple responses to d13C. This d18O plateau was confirmed by repeat variations in mean annual temperature or rainfall. measurements. There is no significant correlation Overall, d13C increased by 2.5x from 1938 to 1992 between d13Candd18O when all the data are (Fig. 7). The beginning of this rise precedes both the considered (r = 0.1; n =44). But, if the 1930–1955 temperature increase and the rainfall decrease by 15 period is considered separately, the year-to-year and 27 yr, respectively. The d18O variations in MND- covariation is quite strong: r =+0.75 (p b0.001; S1 are unrelated to temperature: the positive shift from n =26) across the d18O plateau and +0.92 1930 to 1955 occurs when mean annual temperature is ( p b0.01; n =16) for the data prior to and after the virtually constant, and the increase in d18O since 1970 plateau (Fig. 8), implying that isotopic disequilibrium is in the opposite sense to the isotopic effect of the cannot be ruled out as a cause of minor variation. observed recent warming. Moreover, the ion microp- However, disequilibrium cannot be the cause of the robe data indicate that speleothem d18O is driven high d18O values from 1930 to 1955 because d13Cis primarily by rainfall. The timing of the post-1970 rise P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 27
1911-29, 1956-92 is less subjective in principle, but not appropriate for 1930-55 MND-S1, as the variation within a growth layer is as -7.5 much as 1.5x. With regards to the correlation along r = +0.92 the growth axis, we point out that isotopic disequili- -8.0 brium effects are not dominating the entire MND-S1 ) record since the disagreement between longer-term -8.5 VPDB trends of d13C and d18O suggests underlying fractio- ‰ nation processes that are not common to both. In
C ( -9.0 18
13 MND-S1, a shift in mean rainfall d O by processes δ -9.5 that do not affect soil moisture and cave humidity significantly, such as a reduction in the number of -10.0 r = +0.75 large winter rainfall events, has the potential to affect speleothem d18O but not d13C. Given that the -3.8 -3.7 -3.6 -3.5 -3.4 seasonal variation of speleothem d18O (0.7–1.5x)is δ18 much larger than the longer-term inter-decadal varia- O (‰VPDB) tion (0.3–0.5x), even subtle changes in the amount of Fig. 8. Comparison of conventional analyses of d18O and d13C for winter rainfall could potentially dominate the the d18O plateau (1930–1955) and the periods before and after the d18 18 13 observed changes in mean speleothem O. We plateau (1911–1929, 1956–1992). The d O and d C values for the compare speleothem d18O with the rainfall record in two data clusters show a clear offset, but are otherwise highly correlated. The linear least squares regression equations are detail below. d13C=3.7d18O+3.3 for the plateau and d13C=4.5d18O+8.3 for the periods before and after the plateau. 4.5. Temporal trends in d18O and rainfall
The MND-S1 record for 1911–1992 shows two in d18O is coincident with an increase in speleothem +0.3x shifts, one at 1930 and another post-1970, Mg (Fig. 7) and decrease in speleothem P (not shown) which could be simply explained by changing the [18], which were interpreted previously to be rainfall- frequency of intense events. The 0.3x rise in d18O sensitive trace elements [18]. after 1970 is consistent with the rise in mean rainfall d18O expected with the 20% decrease in regional 4.4. Isotopic disequilibrium processes rainfall which occurred at this time. The rainfall decrease is due to lower May–October rainfall (Fig. A key concern when reconstructing climate using 9) attributed to a reduction in both the frequency speleothem isotope records is determining whether and intensity of large events over the southwest speleothem d18O and/or d13C is affected by disequi- region [37,38]. Mean precipitation-weighted d18O librium processes. The high covariation between these falls from 4.3x to 3.9x if a single large event variables along the growth axis in MND-S1 (Figs. 7 ( 7.81x, 38 mm, 29/09/01) is omitted from the and 8) indicates that kinetic processes may be driving lighthouse record demonstrating the sensitivity of the sub-decadal variability. This is further supported annual rainfall d18O, hence speleothem d18O, to by the slope of the linear regressions (Fig. 8), which these large events. The 1911–1992 lighthouse rain- indicates that the variation in d13C is 3.7 times that of fall record is unreliable, but we adopt the linear d18O during the plateau period, and 4.5 before and regression in Fig. 8 to convert the 1911–1992 daily after. This variation is close to the ratio 3:1 proposed rainfall records to d18O for two high-quality stations by Hendy [11] to indicate isotopic disequilibrium, [20]. We estimate that the rainfall decrease would reflecting the simultaneous but larger effect on d13C raise annual rainfall d18O by approximately 0.1x at versus d18O. Cape Naturaliste (Fig. 1) to 0.4x at Manjimup, with Isotopic disequilibrium is also indicated by a the latter being a minimum value as the lighthouse simultaneous increase of d13C and d18Odowna calibration does not account for overland isotopic growth layer from the speleothem apex [11]. This test distillation effects experienced by non-coastal sta- 28 P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32
200 1911-29 1930-55 due to a reduction in calcite precipitated from May 18 (δ O plateau) to September rainfall relative to the 1956–1970 period because there was less (isotopically light) 150 post-1964 winter rainfall overall (Fig. 9), and speleothem growth rates indicate that calcite is precipitated in narrower bands after 1970 (Fig. 7). On the other 100 hand, we can account for the post-1970 rise in 18 Rainfall (mm) speleothem d O by the reduction of intense winter 50 rainfall events alone. A tendency for higher speleothem d18O during drier periods does not hold for the 1930–1955 interval, as this was the longest wet period on record J F M A M J J A S O N D (Fig. 7). Whether this 0.3–0.5x increase in d18Oin Month the MND-S1 record can be attributed to either a Fig. 9. Average annual cycles of mean monthly rainfall for the change in the frequency of intense winter events or the periods 1911–1929, 1930–1955 (d18O plateau), and 1965–1992. mass of calcite deposited is difficult to assess. 1930–1955 has higher rainfall in June compared with 1911–1929. Averaged monthly rainfall for the periods 1911– This increase in early winter rain during 1930–1955 is due to the higher incidence of meridional troughs, which may bring 18O- 1929 and 1930–1955 (Fig. 9) shows neither an enriched rainfall to the mid-latitudes of southwest Australia. The increase in summer nor a decrease in winter rainfall greatest decrease in rainfall post-1964 is for the winter months during 1930–1955. However, the 1930–1955 period (May–October). shows higher early winter rain (June; Fig. 9) than at any other period on record. Higher early winter rain during this period has been attributed to a higher tions [22]. Moondyne Cave lies between these incidence of meridional troughs [39]. The key link stations on the lee side of the coastal ridge. Thus here may be that meridional troughs are capable of the estimated change in rainfall d18O is approx- transporting tropical moisture, which is isotopically imately equivalent to the rise in speleothem d18O heavy, to the mid-latitudes. after 1970 and the record appears to have captured Early and late winter rain vary slightly in their the rainfall decrease. sources [39]. Southwesterly rain predominates over These arguments assume that the majority of Moondyne Cave in winter, bringing moisture from calcite precipitated in each annual growth increment nearby or higher latitudes. Northwesterly rain is sourced from May to September rainfall. However, derived, in part, from the tropics is more common the 0.7–1.5x annual cycles in speleothem d18O in late summer, but may extend to early winter [39]. revealed by ion microprobe analysis indicate that a Tropical moisture is isotopically heavier than that in significant amount of calcite is also sourced from higher latitudes [22] and may be transported to the isotopically heavier rainfall either from summer mid-latitudes when meridional troughs link tropical events or evaporation of soil water during warmer and mid-latitude low pressure belts, reflecting the months [36]. If this is the case, then small changes in breakdown of the Australian summer monsoon. This the relative masses of speleothem calcite deposited process was observed in the Tasmanian rainfall from isotopically light winter rainfall versus isotopi- d18O dataset when an deast-coast lowT extends cally enriched rainfall/soil moisture could have down the eastern Australian coast, bringing large produced the shifts in mean d18O (~0.3x), deter- amounts of rain with relatively high d18O values mined by bulk sampling of MND-S1, that have a [31]. In the southwest Australian rainfall dataset, complex relation to climate. For example, it is several incidents involving meridional troughs were possible that the proportion of winter speleothem identified for May and June (dashed line, Fig. 4A) calcite could vary as a function of the carbonate from mean sea-level pressure and cloud cover saturation state of these waters. The post-1970 synoptic charts (not shown). But rainfall d18O increase in speleothem d18O may have also been values for these events are not consistently higher P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 29 than those for other events with similar rainfall reaching the speleothem. This suggests that the amounts, indicating that the contribution of a speleothem d13C variations broadly represents the tropical moisture source is minimal. It thus appears cumulative fractions of 13C/12C in the soil carbon pool unlikely that a persistent tropical moisture source above the cave. The lowering of speleothem d13C could have raised rainfall d18O over the period during ~1927–1953 suggests that it is responding to 1930–1955, but a much longer dataset of rainfall the wetter climate at this time. A possible driver of d18O would be needed to confirm this. this ~1x decrease in d13C may be an increase in the 18 A comparison of seasonal d O during the plateau partial pressure of soil CO2 during wetter years if (1930–1955), and before (1955–1970) and after the limestone dissolution is taking place under open- rainfall decrease (1970–1992), may have assisted the system conditions, but this is speculative. interpretation of the higher d18O values seen in the Recent studies have found large (more than several conventional data. But it is highly unlikely that any per mil) fluctuations in speleothem d13C in modern differences would have been detectable outside the [28] and Pleistocene [40] speleothems that were analytical precision. Furthermore, the +0.3–0.5x attributed to an increase in the relative proportion of 18 13 13 step in speleothem d O at 1930 coincides with an C-enriched atmospheric CO2 to C-depleted bio- abrupt change in the calcite fabric from compact genic CO2 in the soil carbon pool. It is possible that inclusion-free closed columnar calcite to inclusion- the thickening of understory C3-type vegetation due rich open columnar fabric (Fig. 2). The open nature to a reduction in regional forest burning may have 13 of the crystallite fabric over this period is not altered the d C of soil CO2. However, if it did, we suitable for ion microprobe analyses as a perfectly would expect to see a post-1970 decrease in speleo- smooth surface is required to reduce topographic them d13C owing to an increase in 13C-depleted effects described earlier. biogenic soil CO2. If the burning itself affected speleothem d13C values, either through vegetation 4.6. Speleothem d13C temporal trends loss or change to the 13C/12C of the decaying matter, then we would expect our conventional d13C analy- Potential climatic effects on speleothem d13C are ses, which typically span 1–3 yr growth, to capture not straightforward because speleothem d13C begins this, but we do not see significant short-term effects. to rise in 1938, 15 yr before temperature begins to Similarly, we cannot attribute the rise in speleothem rise and 27 yr before rainfall decreases. The 2.5x d13C to C4-type plants with metabolic pathways that variation in d13C is too large to be due to the effect favour the uptake of 13C [3]. Grasses make up a minor of the 0.8 8C temperature rise on 13C/12C fractiona- portion of the vegetation covering Moondyne Cave tion [11,16]. A general response to the late 1960s and only 20% of southwest Australian grasses use the decrease in rainfall may be evident if speleothem C4 pathway [41]. If C4 grasses were present within a d13C is divided into two episodes: (1) 1911–1955 thinned understory, they were replaced by thick when speleothem d13C is at or below 9x, and (2) woody C3-type vegetation from about 1980 onwards, 13 12 post-1955 when the largest change in speleothem which would drive soil CO2/ CO2 in the opposite d13C (+1.5x) occurs, peaking at 8x in 1990. sense. Other studies have tracked the reduction in 13 13 12 However, variations in speleothem d C are primar- atmospheric CO2/ CO2 due to fossil fuel burning in 13 12 ily a function of soil CO2/ CO2,andCO2 modern speleothems [42], but this is clearly not degassing rate if the difference in the partial dominating our record. 13 12 pressures of CO2 between the solution and cave Water stress may also raise the C/ C ratio of atmosphere is large [11]. Therefore, we examine all biogenic CO2 by contracting leaf stomata, thus possible impacts of the environmental history on soil restricting CO2 supply [13,14]. A survey of leaf 13 12 13 CO2/ CO2 above Moondyne Cave during the d C in vegetation dominated by Eucalyptus, Acacia, growth of MND-S1. and Melaleuca species along a rainfall gradient in Speleothem d13C values during the 1911–1955 southern Queensland indicated an approximate 1x interval are close to values modelled [11,16] for drip increase in d13C for every 200 mm decrease in rainfall water that has not degassed significantly before [43], consistent with the observed changes in the 30 P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32
MND-S1 speleothem and southwest Australian mete- rainfall d18O signal passes through the cave roof, orological records after the late 1960s. could be a powerful tool for the reconstruction of We consider also that water reduction may also rainfall variability over short timescales. However, bring about kinetic fractionation of 13C/12C in the conventional d13C and d18O analyses of the entire precipitating solution also, argued as follows. During speleothem record (1911–1992) show isotopic trends the drier 1956–1992 interval, the increase in speleo- that cannot be readily reconciled with the instrumental them d13C parallels the increase in speleothem Mg meteorologic record. The lack of simple responses of (Fig. 7), which was interpreted to reflect the longer speleothem d18O and d13C to either climate or non- residence time of water in the overlying limestone due climate factors on decadal timescales calls for inves- to a reduction in rainfall recharge. The primary tigation of another active speleothem from the same mechanism for an increase in the Mg/Ca ratio in area, keyed to the results reported here by using MND-S1 during the prolonged drier period was annual trace element cycles. This is a key aim of interpreted to be prior-calcite precipitation resulting ongoing research and will confirm how robust the in a loss of Ca with continual uptake of Mg, once the response of southwest Australian speleothems is to the water has reached saturation with respect to Ca [44]. rainfall decrease. Drip water and continued rainfall Since this process implies that calcite is precipitated monitoring will enable full examination of the before reaching the speleothem, it also indicates that isotopic signature of meridional troughs. 18 degassing of CO2 from the saturated waters must have For now, the seasonal variation in speleothem d O occurred, which will increase drip water d13C [12]. revealed by ion microprobe analysis indicates that The Pearson’s correlation coefficient between d13C subtle changes in the frequency of intense winter and Mg (when averaged to approximately the same rainfall events and possibly moisture sources could sampling resolution of d13C) is weak (r =+0.29, produce significant changes in mean speleothem d18O. p b0.03), indicating that some other process is The ion microprobe results also raise the possibility contributing, whether this be 13C enrichment of soil that the masses of speleothem calcite deposited in organic matter or isotopic disequilibrium processes. winter and summer could vary as a function of the Isotopic disequilibrium precipitation processes are seasonal drip rate and carbonate saturation state of controlled by the rate of CO2 degassing from the these waters. If this is the case, then small changes in solution, but we cannot identify any event in the the relative masses of calcite deposited in winter and history of Moondyne Cave that could have brought summer could produce significant shifts in mean d18O 13 about a change in either solution CO2 partial pressure and d C that have a complex relation to climate. or cave atmosphere partial pressure at this time, since These findings should be generally applicable to the rainfall, temperature, and vegetation did not alter interpretation of long-term trends in speleothem geo- significantly until after 1964–1970, nor cave use until chemical records for shallow caves where seasonal 1959. Nonetheless, the short-term covariation variations in geochemical tracers are relatively large, between d13C and d18O is very strong, indicating including the sub-tropical monsoon belts and mid to that isotopic disequilibrium processes are important in high latitudes with distinct rainfall seasons. MND-S1, but dominate sub-decadal rather than longer-term variations. Acknowledgements
5. Conclusions We thank Caveworks (Margaret River, WA) for loan of MND-S1 and their assistance, N. Nicholls for In situ ion microprobe analysis of d18O annual the instrumental climate dataset, and S. Eberhard for cycles in stalagmite MND-S1 demonstrates that Moondyne Cave’s history. We also thank Marty seasonal variations of d18O of rainfall are preserved Grove, Joe Cali, Mike Shelley, and Heather Scott- in southwest Australian speleothems. Our findings Gagan for technical assistance; Graham Farquhar and suggest that analysis of annual d18Ocyclesin Hilary Stuart-Williams for water analyses at RSBS; speleothems from shallow sites, where the seasonal Joan Cowley for water d18O analyses at RSES; and P.C. Treble et al. / Earth and Planetary Science Letters 233 (2005) 17–32 31
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