6 Impact of climate variability
compiled into the simulations (due to missing river gauge data). However, Fig. 6.4 shows times series of the Mary River and Burnett River discharge. Two extreme flooding events are visible, one in February 1992 and in February 1999. The latter one lead to significant loss of seagrass in the Great Sandy Strait [Campell and McKenzie, 2004]. Because the peak values of the Burnett River are only about one sixth of the maximum flow of the Mary River, the main focus of the short term variability is on the impact of the Mary River.
Year
River discharge
19900 1993 1996 1999 2002 2005 2008
2000 4000 6000 8000
Mary River Burnett River
Figure 6.4:Freshwater discharge of the Mary and Burnett River (1990-2008) in m3/s. The grey bars indicate El Ni˜no/La Ni˜na events.
6.2.1 Catchment area
The catchment area of the Mary River covers 5000 km2. It reaches from 25.2°to 27°S and from 152° to the coast (see Fig. 2.1), thus a stripe of 150×50 km2. The Mary River flows into the northern region of the Great Sandy Strait draining modified catchment of dryland grazing, agricultural crops, cleared land, forests and both sewered and unsewered urban development areas [Rayment and Neil, 1997]. For an average rainfall year, 21% of rainfall is exported as runoff into the Mary River and 268,000 tonnes of eroded sediments flow into nearshore regions annually. The river further flushes nitrogen (1.7 kg/ha/y) and phosphorus (0.2 kg/ha/y) into the Great Sandy Strait passage each year [Schaffelke, 2002].
6.2 Short term variability
6.2.2 River discharge statistics
The river discharge time series in Fig. 6.4 are rather spiky. The peak values for the 1992 and 1999 flood reached 7000 m3/s. In Fig. 6.5 the cumulative distribution functions (CDF) for the Mary and Burnett Rivers are given. The mean flow for the Mary River is 30 m3/s and 10 m3/s for the Burnett River. The Median is 3 m3/s and 0.8 m3/s, respectively, and therefore only a tenth of the mean. Thus, most of the time both rivers are almost dry. The rare extreme events shift the mean to higher values. The CDF indicates, that the probability to exceed a flowrate of 70 m3/s for the Mary river and 30 m3/s for the Burnett River is less than 5%. Fig.
6.4 further indicate that the high flow volumes are strongly linked to El Ni˜no/La Ni˜na events.
1 10 30 100 1000
0.01 0.1 0.5 1
River discharge
CDF
Mary River Burnett River Median
Figure 6.5: Cumulative distribution functions (CDF) of the freshwater discharge (in m3/s) of the Mary and Burnett River (1990-2008).The two dashed lines indicate the mean.
6.2.3 Flooding events
In Fig. 6.6 the response of Hevey Bay to flooding events is shown. Plotted are the depth aver-aged salinity field and a transect in the southern part of the bay, 10 days past the peak flow of the Mary River. The river discharge associated with these flooding events is 7100 m3/s (1992), 6700 m3/s (1999) and 900 m3/s (2008), see also Tab. 6.2. The peak value of 900 m3/s seems rather low, compared to the 1992 event. However, the 2008 flood was preconditioned by three 750 m3/s peaks (in the 40 days before the flood) and puts it therefore in a comparable range to the 1992 event. Fig. 6.6 indicates, that the outflow of the freshwater is restricted to a narrow region along the western coast. The transects for 1992 and 1999 further indicate a pronounced frontal structure (horizontal and vertical). Beside this narrow coastal freshwater strip, the whole bay is mainly unaffected by the flood. Although the river discharges for the 1992 and
6 Impact of climate variability
Table 6.2: Atmospheric condition and river discharge for three flood events.
Year River discharge Wind direction Wind speed
1992 7100 m3/s SE 6 m/s
1999 6700 m3/s S 8 m/s
2008 900 m3/s SE 3 m/s
1999 flood are comparable, the transects in Fig. 6.6 show a different behaviour. Whereas in 1992 the bay has a nearly uniform salinity distribution, strong salinity stratifications are visible for 1999. These differences are mainly caused by the location of the strong rainfalls. For 1992, they occurred mostly in the southern parts of the Mary River catchments combined with minor precipitation in Hervey Bay. For 1999, the rainfalls were uniformly distributed over Hervey Bay and the catchments. Thus, due to the heavy precipitation, a freshening of the surface layer is visible (Fig. 6.6b) and thus explaining the vertical salinity stratification. This further explains the greater width of the river plume. Both events are assisted by strong southerly winds. Hence, a northward flow-through in the Great Sandy Strait prevents an outflow of the Mary River discharge into the southern region of the Strait. This high flow-through further pushes the riverine water quite effectively into the bay. This is not the case for the 2008 floods.
Due to the light winds, the fresh water remains in the northern part of the Strait. Moreover, the saline water in the bay acts as a salt barrier, which prevents the transport of the riverine fresh water into Hervey Bay. Further, the low river discharge leads only to a weak coastal plume on the western shore (Fig. 6.6c).
(a)
−25.2
−24.8
(b)
−25.2
−24.8
(c)
152.6 153
−25.2
−24.8
35 35.2 35.4 35.6 35.8 36 36.2 36.4 36.6
−15
−5
−15
−5
152.6 153.0
−15
−5
Figure 6.6: Depth averaged salinity (in psu) for a) 4 Mar. 1992, b) 20 Feb. 1999 and c) 23 Mar.
2008. In the left column the salinity transects along the red lines are shown.
6.2 Short term variability
6.2.4 The flood of 1999
In the following, the 1999 flood event (February 1999) [Campell and McKenzie, 2004] is used to estimate typical exchange time scales associated with high riverine flow, and also a time which is needed for Hervey Bay to recover to a “normal state”.
Two experiments are conducted. In the first one, called the flood run (FR), the river discharge, atmospheric boundary conditions and open ocean boundary conditions are prescribed using the forcing given in Sec. 2.2. At the same time, a neutral tracer with a concentration of 100 units was released into the river. In the second experiment, the control run (CR), the high river discharge due to the flood is completely switched off. Both experiments start at the 1.
February 1999 and run for three months. Fig. 6.7 shows the impact of the 1999 flood event.
d e f a)
152.6 153.0
−25.2
−24.8
0 2000 4000 6000
River discharge
b)
Feb Mar Apr May
−0.03
−0.02
−0.01 0
∂ S
c)
20 25 30 35
Salinity
d)
10 20 30 40
Tracer
30 35
Salinity
e)
0 10 20
Tracer
30 35
Salinity
f)
Feb Mar Apr May0
20
Tracer
Figure 6.7: a) position of three virtual measurement stations to showing the impact of the 1999 flood, b) river discharge in m3/s, c) salinity gradient ∂S in psu/km (see also Fig. 5.8). The blue curve represents the flood run (FR) and the red curve the control run (CR). Pictures d-f shows the salinity at the three stations d-f. The green curve indicates the tracer concentration.
The peak river discharge is approx. 7600 m3/s at the afternoon of the 9 February. In the beginning of March a second minor flood occurred with a peak flow of 800 m3/s. In Fig. 6.7c the impact of the flood on the salinity gradient is visible. The transect is positioned close to station f. The minimum gradient has a delay of 6 days compared to the flood peak. This corresponds to a plume velocity of 10 cm/s. Taking the wind conditions for this event into account (Tab. 6.2), clearly shows, that the plume is advected along the western shore, due to the wind induced currents (Fig. 4.2). Fig Fig. 6.7c further indicates that, although the salinity gradient shows a significant minimum, Hervey Bay completely recovers to the undisturbed state (CR) within two months.
6 Impact of climate variability
Fig. d-f depicts the change in salinity and tracer concentration at the three stations. A closer inspection of the time series yields, that the minimum in the salinity and the maximum in concentration appear simultaneously. Thus, both are advected with the same velocity. More-over the further north the station is situated, the weaker the flood impact. This seems quite reasonable, because the freshwater plume is much longer exposed to entrainment and tidal mixing. Nevertheless, the salinity time series show the same recovery time of 2 months to the undisturbed state.
5 10 15 20 25 30 35 40 45 50
−1
−0.5 0 0.5 1 1.5 2 2.5 3 3.5
Days past peak flow
log(salinity difference)
Point d) Point e) Point f)
Figure 6.8: Time series of the logarithmic salinity difference (CR-FR) for the three virtual measure-ment stations (see Fig. 6.7a). The dashed lines indicate a linear fit. For visualisation, the time series are shifted along they-axes
To compute a second time scale, it is assumed that the salinity time series of experiment FR recovers exponentially to experiment CR. Fig. 6.8 depicts the logarithmic salinity difference between the experiment FR and experiment CR. The linear fits indicate that the exponential recovery is a valid assumption. Further, the exponential decay is nearly the same for all three stations. The linear fits indicate a decay constant of approx. 20-24 days. This time scale is similar to the flushing time of the western part of Hervey Bay under SE wind (Tab. 4.3). Thus, the flushing of the riverine freshwater is strongly affected by wind conditions at the time.
6.2.5 Flood response
In the previous section a detailed analysis of the 1999 flood was given. The same exercise was repeated with the events of 1992 and 2008. In Tab. 6.3 the results are summarised. The
6.2 Short term variability
simulation indicates, that although the flood related river discharges differs significantly, the recovery times vary only slightly. For all three events, Hervey Bay shows a decay rate of the disturbance, of approx. 22 days. This recovery time scale to the experiment CR is seen in the salinity and in the salinity gradient. The simulations further indicate, that the flood response is closely related to the wind induced residual circulations. The exponential decay with approx.
22 days, can also be seen in the residence time for SE wind (Tab. 4.3).
Table 6.3: Exponential recovery time for the three flood events in days. The salinity recovery is averaged over the three stations. The salinity gradient∂S is computed along the transect indicated in Fig. 3.1. For 2008, this measure could not be computed.
Year River discharge salinity recovery ∂S recovery
1992 7100 m3/s 26 d 20 d
1999 6700 m3/s 22 d 18 d
2008 900 m3/s 20 d
-7 Gravity currents
7.1 Release of gravity currents
The density gradient time series in Fig. 5.8a show, that the density at the shore is higher than on the shelf. This leads to the establishment of an inverse circulation. Because the positive density gradient is a gravitational unstable state and therefore gravity induced flows are triggered. This flow of dense water originating from cooling, evaporation, or salinisation on the shelf, spills over the shelf edge and can develop as near-bottom gravity current or an intermediate-depth intrusion. Quite often, it is difficult to observe them in nature due to their intermittent character. It is worth mentioning that until now no observations of the gravity flows in Hervey Bay are available. The main research focus was on Hervey Bay itself; therefore, no field measurements were taken on the northern shelf. Middleton et al.
[1994] lacked observational evidence in support of their hypothesis that Hervey Bay potentially exports high salinity water formed through a combination of heat loss, high evaporation, and weak freshwater input in shallow regions of the bay. ... Thus, a second hypothesis is that the high nutrient, low-oxygen waters that constitute the anomalous water masses observed at both the Sandy Cape and Double Island Point sections consist partly of cooler, saltier ’winter mangrove waters’ exported both north and south of the Great Sandy Strait on each flood and ebb tide. The exported waters would subsequently sink off the continental shelf to their own density level, progressively mixing with ambient waters... Thus the common assumption was, that tidal flushing would lead to constant export of this dense water to the continental shelf. Further due to its low aspect ratio and therefore the lack in supporting high density gradients, significant gravity currents ( [Tomczak, 1985; Lennon et al. , 1987; de Silva Samarasinghe, 1998] were not expected.
Fig. 7.1a shows such a flow event in June-July 1995. The density within Hervey Bay reaches values of greater thanσt=26.2 kg/m3, which is equivalent to a depth of approx. 300 m. During its way down the shelf, the plume is channelled in the Mary River Canyon, to finally reach a depth of 200 m.
7.1.1 Formation
To describe the four different stages in the development of the gravity flows, the classification
7.1 Release of gravity currents
a)
26.2 25.2 25
152 152.5 153
−25
−24.8
−24.6
−24.4
−24.2
−24
17 20 b)
152 152.5 153
−25
−24.8
−24.6
−24.4
−24.2
−24
c)
35.6 36
152 152.5 153
−25
−24.8
−24.6
−24.4
−24.2
−24
Distance [km]
Depth [m]
d)
0 50 100
−200
−150
−100
−50 0
24.5 25 25.5 26
Figure 7.1: a)bottom densityσt[kg/m3] on the 28 July 1995, b) bottom temperature - [C], c) bottom salinity - [psu] and d)σ-density along the black line in b) [kg/m3]. The thick dashed lines in a,b and c are the 40 m, 100 m and 300 m depth isolines.
proposed in Shapiro and Hill [2003] is used. The pre-conditioning is the stage when dense water accumulates on the shelf and a density front is formed. The short active stage corre-sponds to the period when the leading edge of the dense water accelerates down-slope. The main stage relates to a quasi-steady flow with a noticeable down-slope component. These two stages can also be combined in a down-slope propagation stage. The final stage is reached when the water spreads isopycnically off the slope, but traces of the cascade may still be detected by inclined isopycnals over the slope.
7.1.2 Pre-conditioning
To trigger gravity currents in Hervey Bay, three conditions are necessary. First, the density gradient has to exceed 0.008 kg/m3/km. This is a rather moderate gradient and the same as the threshold to define inverse circulations. The second condition, which is necessary, is that the SST in Hervey Bay has to drop below 20°C. This condition is mostly fulfilled during June/July. The salinity gradient is of minor importance. The dependence on the SST indicates, that Hervey Bay has a temperature driven cascade [Shapiro and Hill, 2003], with a response
7 Gravity currents
to surface cooling assisted by advection of salinity. The third mechanism is the tidal forcing.
The initialisation of the plume coincides with the occurrence of neap tide. During this time tidal mixing is sufficiently reduced and a two layered flow structure can develop.
Wind forcing is not directly involved into the triggering of the plumes. It is only important to maintain the gradients across the shelf. The wind forcing is important to restrict the path of the plume between Breaksea Spit and Lady Elliot Island (see Fig. 3.1). Such an event is shown in Fig. 7.1. In the northwestern part of the shelf, water of low density and higher temperature is situated. This acts as an effective barrier for a northward flow of the plume.
Hence, the whole flow is directed into the Mary River Canyon. This is a typical situation during northeasterly wind. Relative warm water from the southern part of the Great Barrier Reef is pushed southwards onto the northern shelf of Hervey Bay. Further, the saline water from the northern shelf is advected into the bay. This mechanism supports the intensification of the flow but is not necessary to actually release the plume.
7.1.3 Down-slope propagation
The flow of the plume is mainly controlled by the steepest decent into the Mary River Canyon.
Therefore, topography suppresses the effects of earth rotation until the plume reaches the 100 m depth isoline. Although the flow is disturbed by the bathymetry, a good estimate to compute the frontal velocityuN of is given by Nof [1983].
uN of =g′tanα
f with g′ = ∆ρ
ρ0 (7.1)
where g′ is the reduced gravity within the plume, f is the Coriolis frequency and α is the bottom slope. To compute this theoretical velocity f is taken to be 3·10−5 s−1 and α is estimated with 0.0016. To compute the reduced gravity, the density gradient (see Fig. 5.8a) is used. The transect to calculate the gradient is aligned with the flow path of the plume and has a length of 60 km. Using a density gradient of 0.008 kg/m3/km, equivalent to a density difference of 0.5 kg/m3, leads to a maximum frontal speed of 0.25 m/s. This velocity is also seen in the simulations. The average flow velocity in the steady state is approx. 5-7 cm/s.
Inspection of the time series shows that the plume front passes the 100 m isolines approx. 5-8 days past neap tide. This can be explained by simple geometry. The gravity current needs on average 6 days to travel 150 km, which is approx. the shelf width. Therefore, the reduced mixing during neap tide is necessary to trigger the release of the plumes.
If the plume has passed the 60 m isoline, the slope of the shelf nearly doubles (see Fig. 7.1d).
According to Eq. (7.1) this would cause a doubling of the frontal speed, which is not seen in the simulations. Due to dispersion and intrusion, the density gradient across the plumes starts to weaken, reduces therefore g′ and balances thus the change in bottom slope.
After passing the 60-80 m isoline, the plume is no longer restricted in its flow path by the